The term “maritime continent” was first coined by Ramage (1968) to describe the region including the Indonesian Archipelago, tropical Northern Australia and Papua New Guinea, emphasizing the important role that the numerous small islands played in the local weather. This region of the tropics has also come to be known as the “boiler box”. Through a combination of shallow ocean depths and relatively high solar insolation , the warmest average sea surface temperatures over the whole globe occur in this region, representing the large amount of energy being stored in the tropical western Pacific Ocean. Figure 1 shows an annual averaged sea surface temperature for the entire globe. It is clear that the warmest temperatures, 30oC or higher, are located around the maritime continent region. The warm sea surface temperatures, coupled with the diurnal heating cycle, local topography and the prevailing synoptic flow dominate the convective activity over this region. To quantify the extent of this convection compared to the rest of the globe, figure 2 shows the mean outgoing longwave radiation for the period from 1979-1995. This shows outgoing longwave radiation as low as 205 W/m2 which corresponds to a brightness temperature of 245.2K (= -27.9oC) . These low averages of outgoing longwave radiation are indicative of the cold cloud tops of the deep tropical convection in this region. The tropopause in this region is high and cold, such that convective cells can penetrate to heights greater than 20km and pressures less than 100mb.
A significant fraction of the total rainfall throughout the maritime continent is produced by thunderstorms with the particular feature of remaining fixed over the islands throughout their development phase (Keenan et al., 1989). Figure 3 (Holland and Keenan, 1980) shows a typical example of the evolution of clouds over the maritime continent region during the course of the middle part of the day. Holland and Keenan (1980) point out that the convective clouds in the afternoon are linked to the topography, such that they virtually provide a mapping of the region. Every island and mountain is marked by the presence of its own cumulonimbus. Despite the presence of larger scale oceanic disturbances, and regardless of the phase of the Madden-Julian oscillation these locally generated, diurnally forced thunderstorms are the basic convective element in this region (Keenan et al., 1990). Johnson and Priegnitz (1981) emphasize that the diurnal forcing of convection is important in this region under both synoptically disturbed and undisturbed conditions. These storms are the deepest convective activity observed on a regular basis anywhere in the entire world.
The resulting redistribution of heating due to the convective elements is an important factor in the dynamics of the tropics, transfers between the troposphere and stratosphere, and global teleconnections. For this reason, there have been a great deal of observational studies in the tropical western Pacific in order to try to gain an understanding of the characteristics of convection in this region. These include the Winter Monsoon Experiment (WMONEX) (Houze et al., 1981), the Australian Monsoon Experiment (AMEX) (Holland et al., 1986), Stratosphere - Troposphere Exchange Project (STEP)(Russell et al.,1993), and the Equatorial Meso-scale Experiment (EMEX). More recently, the Island Thunderstorm Experiment (ITEX) (Keenan et al., 1989) and the Maritime Continent Thunderstorm Experiment (MCTEX)(Keenan et al.,1996) have concentrated on observations of the diurnally forced island thunderstorms. As a comparison to the climatological values, during STEP, Knollenberg (1993) made measurements of cloud top temperatures from the ER-2 aircraft. He recorded cloud top heights of 17 km in the anvil shield of towering cumulonimbus with cloud top temperatures of -90oC. Satellite estimates of the equivalent radiative temperature showed a much higher temperature of -60oC. This would suggest that this was an optically thin section of the anvil, and we would expect that the convective turret penetrated to much higher heights, an indication of the intensity of the convection in this region.
This review paper will investigate the role of the maritime continent convection in the global circulation. The importance of this region in the maintenance of the Hadley and Walker circulations has been suggested by several authors (e.g. Bjerknes,1966;Ramage,1968). It follows that the convection in the maritime continent region is significant in the global context.
In 1966, Bjerknes investigated the response of the Hadley circulation to anomalies of ocean temperatures along the equator. He reasons that a warmer than normal equatorial ocean over a wide span of longitude will make for a stronger Hadley cell, in that it will circulate faster and will transport angular momentum to the subtropical jet at a faster rate. Bjerknes (1966) studies this through a comparison of the winters of 1957 and 1958. He shows that for the winter of 1957-1958, an ENSO year, when the warmest water in the eastern Pacific is displaced farther south along the equator, the Hadley circulation is stronger, transporting more angular momentum northward which is manifested in increased westerlies in mid- latitudes. If we make the connection that the warmer sea surface temperatures correspond to the case of prominent convection in the maritime continent, we see that the work of Bjerknes (1966) emphasizes the important role of the maritime continent region on the Hadley circulation. Figure 4 shows some data supporting Bjerknes theory. The average distribution of atmospheric pressure at sea level is shown for three separate winters. 1957-1958 was an ENSO event, and thus anomolously warm surface ocean waters occured along the equator of the Pacific Ocean. For comparison, two non-ENSO years are also shown. As predicted by the theory of Bjerknes (1966) during the ENSO event the subtropical high is displaced further to the south in the eastern Pacific, compared to non-ENSO years. This is in accordance with the southern displacement of the warmest surface waters in this region from north of the equator to directly on the equator (Bjerknes, 1966).
In the classic paper by Ramage (1968), a comparison study is presented of two very different Januarys over the western Pacific. In January 1963, the jet stream over the western Pacific was much stronger and displaced farther to the south than usual. The Siberian Low and Aleutian Low were more intense than usual, and the rainfall was above normal over the maritime continent (Ramage, 1968). For the January 1964 case, meridional flow dominated the eastern hemisphere with the subtropical jet located in its average position. The Aleutian Low was located further east than it normal position, and the pressure gradients to the Siberian Low were only half the values for the January 1963 case. Rainfall was below normal over most of the maritime continent (Ramage, 1968). Figure 5 shows the change in monthly rainfall from January 1963 to January 1964. If we consider these two cases to be extremes within which the typical western Pacific weather conditions will occur, then Ramage’s (1968) comparison of these cases should represent a range of effects of the conditions in the tropical western Pacific on the larger scale circulation.
In order to understand the details of the interactions between the tropics and mid-latitudes, Ramage (1968) looked specifically at the 200 mb kinetic energy, temperatures and pressures. Figure 6 shows the ratio of 200 mb kinetic energy of January 1963 to January 1964 and Figure 7 shows the change in 200 mb mean height from January 1963 to January 1964. The major points to be taken from these 2 figures is that the circulations were much stronger in 1963 and the 200 mb temperatures and 200 mb heights were greater in 1964. This presents us with the problem that the largest positive differences in 200 mb height, correspond to the largest negative differences in latent heat release, i.e. rainfall. This suggests that there must be some different or enhanced mechanism operating during January 1963 in order to remove heat from the tropical region. Ramage (1968) suggests that the meridional circulation in 1963 has enhanced surface northerlies which caused more water to be evaporated from the ocean maintaining the heavy rains over the maritime continent. The energy derived from the rain then helps the circulation keep its strength. The heat source in the tropics and the heat sink south of the mid-latitude jet are connected by a vigorous meridional circulation. Therefore the heat that is brought to the upper troposphere by the maritime convection has an outlet and is transported to the north through the Hadley circulation avoiding the buildup of heat in the tropical upper troposphere. This keeps the atmosphere in the region unstable and favorable for more convection. Hence, there is a positive feedback such that the more vigorous circulation leads to enhanced convection which further drives the circulation.
In the January 1964 case, the subtropical jet was farther north, and therefore transport of heat from the tropics to the north was restricted. Thus despite less rainfall over the maritime continent heat was unable to accumulate in the upper tropical troposphere, and convective instability was reduced. This weakened the Hadley circulation, which caused weakened surface northerlies and therefore less evaporation at the surface further aiding in the inhibition of the convection. The circulation probably gained more of an eastward component in its circulation with rising motions concentrated over New Guinea and sinking motions to the northwest over the Indonesian continent.
This comparative study by Ramage therefore emphasizes the critical feedback system involving the convection in the tropical western pacific, the subtropical jet and the extent of the Hadley cell. Finally Ramage (1968) suggests that if we compare the upper level circulation of the maritime continent region with other parts of the tropics, we will find that the wintertime circulation is strongest in this region suggesting that this feedback with the Hadley cell will be most pronounced in this area. This work is consistent with the ideas of Bjerknes (1966).
The convective activity in the tropical western Pacific is not only a driving force in the north-south Hadley Circulation, but also drives the east-west transport known as the Walker Circulation. Figure 8 shows a schematic of the Walker Circulation during non-ENSO conditions (Peixoto and Oort, 1992 ). The strong upward motions over the maritime continent region drive the upward branch of this circulation which then travels to the east descending over the cold waters of the eastern equatorial Pacific. The circuit is completed with surface easterly flow. Another cell of this east-west circulation is also fed by the upward motion in the maritime continent with its descending branch over the cooler western Indian Ocean.
The role of the maritime continent convection in the Walker circulation suggests that the convective activity in this region will also be related to the occurrence of El Ni¤o-Southern Oscillation (ENSO) variations. Philander (1983) investigates the climatic variations for 7 ENSO events that have occurred since 1950. He describes 3 distinct phases in a “typical” ENSO event: Precursors, growth of anomalous conditions and return to normal conditions. The maritime continent convection mainly manifests itself in Philander’s precursor stage. One of the precursors of ENSO is the eastward displacement of the upward branch of the Walker Circulation to the region between New Guinea and the dateline (Philander, 1983). This would suggest that there is a suppression of the convective activity over the Indonesian Archipelago, and we would expect to see signatures in the maritime continent region such as , decreased precipitation, decreased cloudiness and increased surface pressure. As an illustration of this, Figure 9 shows a time series of the 30-day surface pressure anomalies at Darwin, representing the tropical western Pacific, and Tahiti, representing the tropical central Pacific for the current (1997) ENSO event. Here we see the signatures described by Philander (1983). The Darwin station shows surface pressures higher than normal, representing the movement of the upward branch of the Walker circulation to the east and decreased mean vertical motions over the western Pacific. Likewise in the central Pacific, at Tahiti, the pressures are , in general, lower than normal as the upward branch has moved over this region. Similar signatures would be seen in plots of the rainfall and outgoing longwave radiation anomalies.
To further investigate the structure and variation of the Walker Circulation, Julian and Chervin (1978), looked at wind data from several upper air stations along the equator. They looked at several intervals over the last 20 years, and divided them into periods of maximum and minimum SO Index, and then did a composite for each case. There results are shown in Figure 10. Four upper air stations were used in this analysis, Cocos Island (C),Port Darwin(D),Canton Island(Ca) and Lima, Peru(L). The top half of the figure show the Walker Circulation under “normal” non-ENSO conditions, with rising motions and relatively lower pressures over the maritime continent and subsidence over the eastern equatorial Pacific. The bottom half of the figure shows the Walker Circulation in ENSO years when there are anomalously warm waters over the eastern equatorial Pacific. In this case, Julian and Chervin (1978) show that the area of rising motions has moved over the Central Pacific, and over the Maritime Continent there is actually subsidence.
The convection in the maritime continent is intimately tied to the zonal Walker circulation as a major driving mechanism, and plays a role on the longer time scale variations in that circulation associated with ENSO events.
Brewer (1949) suggested a mechanism by which air is transported to the stratosphere across the cold tropical tropopause by a slow mean motion where it is freeze-dried. Kley et al. (1979) measured stratospheric water vapor mixing ratios of just under 3 ppmv over Brazil. One can calculate that in order to reach this concentration, the air must have been cooled to a temperature of -84oC , assuming saturation with respect to ice. Kley et al. (1979) determined that the local temperature of the tropical tropopause was only -80oC, leading to the suggestion that this air must have entered the stratosphere at some other location.
Newell and Gould-Stewart (1981) analyzed worldwide 100mb temperatures (as an approximation of the tropopause temperature) over a period of seven years, in order to try to identify the region of cross- tropopause transport. They identify areas where tropopause temperatures are sufficiently cold as to dehydrate the air using a criterion of a tropopause temperature of <-82.4oC, which corresponds to an ice saturation value of 3.5 ppmv. This is a slightly higher mixing ratio than the measurements made by Kley et al. (1979) and may therefore result in an overestimation of the entry area. Figure 11 shows a map from Newell and Gould-Stewart (1981) identifying the region where temperatures <-82.4oC were observed at 100 mb, for the month of January. This figure shows that the largest region meeting the temperature criterion was in the Western Pacific in the December - January season. This suggests that rather than the zonally constant transport at the equator proposed by Brewer (1949) there must be a favored region of cross- tropopause transport located over the tropical Western Pacific. Newell and Gould-Stewart (1981) termed this restricted region of transport into the stratosphere, the “stratospheric fountain.”
Further evidence supporting the idea of a restricted region of cross-tropopause transport are shown in Figure 12 (Danielson, 1982). This figure compares temperature profiles of Darwin, Australia (in the stratospheric fountain region) and Panama. The tropopause temperatures at Darwin are sufficiently cold to dehydrate the air, whereas at Panama, the minimum mixing ratio would only reach 8.2 ppmv (again assuming ice saturation).
There is a general agreement on where the bulk of the tropospheric air enters the stratosphere, but the mechanism by which this transport takes place is still an area of much debate. Differing mechanisms have been proposed by Newell and Gould-Stewart (1981) and Danielson (1982).
Newell and Gould-Stewart (1981) define the “stratospheric fountain” area and then go on to describe the transport through this region by mean mass flux calculations. They take previous estimates of zonally averaged upwelling and confine the corresponding mass flux to the fountain region. These calculations yield an average vertical motion of 0.5x10-2 m/s through the entire region. They suggest that this mean large-scale ascent is responsible for the transport and subsequent dehydration of air entering the stratosphere.
The tropical Western Pacific is an area of very strong convection (Ramage, 1968). Danielson (1982) suggests that the strong convection in this area must play a role in the cross-tropopause transport and therefore the dehydration mechanism. Consider a parcel of air rising moist adiabatically in a cumulonimbus turret. The turret will rise with positive buoyancy until it reaches the tropopause. At some height just above the tropopause, this parcel will reach its equilibrium height. However, because of its kinetic energy, the parcel will overshoot this equilibrium level. As the turret decelerates it will entrain warm, dry stratospheric air. This mixing with stratospheric air will alter the equilibrium level of the parcel to be in the lower stratosphere. The turret collapses and spreads horizontally, forming a large cirrus anvil composed of ice crystals. The anvil maintains static instability and buoyancy generated turbulence through radiative warming at anvil base due to upward radiation emitted by the warm surface, coupled with radiative cooling at cloud top. This instability will sustain mixing through the anvil layer resulting in an adiabatic temperature profile. Figure 13 shows a representation of this anvil formation.
The result of this radiative heating profile will be an upward turbulent flux of heat and water vapor. As water vapor and ice crystals are transported upward, cooling will produce supersaturation and thus, the larger ice crystals will grow until they are large enough to fall out. As larger crystals fall out of the volume, the smaller crystals will then have a chance to grow, until they are large enough to fall out. Thus, the downward flux of ice crystals act to remove water vapor from the air being transported upward in the anvil region. The result is that the air just above the anvil layer will be dried to the ice saturation values of the local temperatures. This could in principle continue until observed stratospheric water vapor contents are reached.
This review paper has described the role of the convection over the maritime continent in the global circulation by investigating its influence on global scale meridional circulations, zonal circulations and vertical transports in the atmosphere. The major role played by this region has led to a great deal of observational, theoretical and modeling studies of this region in recent years and this trend should be expected to continue as it seems we must understand the “boiler box” if we are to understand the global circulation.
