Variability of the Caribbean Low-Level Jet and its relations to climate
A maximum of easterly zonal wind at 925 hPa in the Caribbean region is called the Caribbean Low-Level Jet (CLLJ). Observations show that the easterly CLLJ varies semi-annually, with two maxima in the summer and winter and two minima in the fall and spring. Associated with the summertime strong CLLJ are a maximum of sea level pressure (SLP), a relative minimum of rainfall (the mid-summer drought), and a minimum of tropical cyclogenesis in July in the Caribbean Sea. It is found that both the meridional gradients of sea surface temperature (SST) and SLP show a semi-annual feature, consistent with the semi-annual variation of the CLLJ. The CLLJ anomalies vary with the Caribbean SLP anomalies that are connected to the variation of the North Atlantic Subtropical High (NASH). In association with the cold (warm) Caribbean SST anomalies, the atmosphere shows the high (low) SLP anomalies near the Caribbean region that are consistent with the anomalously strong (weak) easterly CLLJ. The CLLJ is also remotely related to the SST anomalies in the Pacific and Atlantic, reflecting that these SST variations affect the NASH. During the winter, warm (cold) SST anomalies in the tropical Pacific correspond to a weak (strong) easterly CLLJ. However, this relationship is reversed during the summer. This is because the effects of ENSO on the NASH are opposite during the winter and summer. The CLLJ varies in phase with the North Atlantic Oscillation (NAO) since a strong (weak) NASH is associated with a strengthening (weakening) of both the CLLJ and the NAO. The CLLJ is positively correlated with the 925-hPa meridional wind anomalies from the ocean to the United States via the Gulf of Mexico. Thus, the CLLJ and the meridional wind carry moisture from the ocean to the central United States, usually resulting in an opposite (or dipole) rainfall pattern in the tropical North Atlantic Ocean and Atlantic warm pool versus the central United States.
The Caribbean Sea is bounded to the south by South America, to the west by Central America, and the north by the Greater Antilles (Cuba, Haiti, Dominican Republic, and Puerto Rico). It is connected to the Gulf Mexico by the Yucatan Channel at its northwest side and is linked with the tropical North Atlantic Ocean in the east. As part of the Western Hemisphere warm pool (Wang and Enfield 2001, 2003), the Caribbean Sea features a body of very warm water (warmer than 28.5°C) during the summer and fall (season always refers to the boreal season). Variability of the warm pool can affect the summer climate of the Western Hemisphere (Wang et al. 2006, 2007). Additionally, tropical cyclones can be formed in the Caribbean Sea or be intensified when they pass over Caribbean warm water (e.g., Shay et al. 2000). Thus, the Caribbean Sea is an important region for both weather and climate.
The present paper analyzes some of available data and has two major purposes. First, it shows in detail the seasonal and anomalous variations of the CLLJ, and discusses CLLJ’s variability from the perspective of ocean–atmosphere interaction. Second, the paper shows and documents the CLLJ’s relationships with the climate of the northern Western Hemisphere. The paper is organized as follows. Section 2 describes the data sets and methods that are used in this study. Section 3 shows the seasonal variation of the CLLJ. Section 4 documents variability of the CLLJ on interannual and longer timescales and shows CLLJ’s relations to the climate. Finally, Sect. 5 provides a summary.
2 Data sets and methods
Many data sets are used in this study. The first one is the National Centers for Environmental Prediction–National Centers for Atmospheric Research (NCEP–NCAR) reanalysis field on a 2.5° latitude by 2.5° longitude grid (Kalnay et al. 1996). Variables that we analyze in this study include monthly sea level pressure (SLP) and horizontal wind velocity at 925, 850, and 200 hPa from January 1950 to August 2006. The second data set is an improved extended reconstructed monthly sea surface temperature (SST) data set on a 2° latitude by 2° longitude grid beginning January 1854 (Smith and Reynolds 2004), but here we only analyze monthly SST from January 1950 to August 2006 for consistency with the data record length of the NCEP–NCAR reanalysis. Another data set is monthly precipitation product of the Global Precipitation Climatology Project (GPCP) (Adler et al. 2003) that is similar to the CPC (Climate Prediction Center) Merged Analysis of Precipitation (CMAP) (Xie and Arkin 1997). The GPCP data set blends satellite estimates and rain gauge data on a 2.5° latitude by 2.5° longitude grid from January 1979 to August 2006. With the data sets, we first calculate monthly climatologies based on the full record period and then anomalies are obtained by subtracting the monthly climatologies for each data set from the data. Our analyses include the calculations of indices and linear correlations.
3 Seasonal variability
Previous studies have shown that the CLLJ has a maximum of easterly zonal wind in the lower troposphere around 925 hPa (Amador 1998; Amador and Magana 1999; Poveda and Mesa 1999; Mo et al. 2005; Poveda et al. 2006; Wang and Lee 2007; also see Fig. 1a, c). The 925-hPa zonal winds from the NCEP–NCAR reanalysis during July and January are shown in Fig. 1b, d. A longitudinal band of strong easterly zonal wind [induced by the North Atlantic Subtropical High (NASH)] is located in the tropical North Atlantic along 15°N. As the easterly trade wind continues to flow westward to the Caribbean Sea, it intensifies forming the CLLJ with the easterly wind larger than 13 m/s in the summer. Figure 1b, d also show that after the easterly wind passes the Caribbean Sea, the easterly wind contours have two relative maxima: one turning toward the western Gulf of Mexico and the United States, and the other one continuing westward across Central America. The former corresponds to the Great Plains Low-Level Jet (GPLLJ) that transports moisture northward for rainfall over the central United States, whereas the latter carries the moisture westward to Central America and the eastern North Pacific (e.g., Wang et al. 2007).
4 Interannual and longer timescale variability
In this section, we first define an index for measuring the anomalous CLLJ and show CLLJ’s anomalous variability. We then examine CLLJ’s relations to the climate in the northern Western Hemisphere.
4.1 CLLJ Index
4.2 Relation to SLP anomalies
In the winter, the extratropical North Atlantic shows a negative SLP correlation and the North Pacific has a positive correlation (Fig. 8a). We keep in mind that in the winter the North Atlantic and Pacific are under the influence of the Icelandic low and the Aleutian low, respectively. The North Atlantic Oscillation (NAO), represented by the meridional SLP seesaw between the NASH and the Icelandic low, links the SLP in the subtropical Atlantic with that in the North Atlantic. A strengthening of the NASH, usually associated with strengthening of the easterly CLLJ, thus tends to correspond to a strengthening of the Icelandic low (with negative SLP anomalies). This explains the negative correlation between the CLLJ and the SLP anomalies in the North Atlantic. The positive SLP correlation in the North Pacific shows that the easterly CLLJ is associated with a weakening of the wintertime Aleutian low (with positive SLP anomalies) in the North Pacific.
During the summer (Fig. 8b), an additional positive SLP correlation is centered in the southwest coast of North America–the North American monsoon region. This shows that a strong (weak) summertime easterly CLLJ corresponds to positive (negative) SLP anomalies near the North American monsoon region and thus a weak (strong) summer monsoon. This is again consistent with the modeling result of Wang et al. (2007) who show that the effect of the Atlantic warm pool is to weaken the CLLJ and to decrease SLP northwest of the warm pool. The physical basis for this response is Gill’s (1980) theory that shows a Rossby wave (with low SLP) to the northwest of atmospheric heating. Figure 8b also shows a negative SLP correlation in the subtropical North Pacific, indicating that the summertime strength of the North Pacific subtropical high is inversely related to the CLLJ.
4.3 Relation to SST anomalies
The difference between the winter and summer Atlantic SST correlation is that in the winter it displays an alternating tripole pattern of zonally oriented negative-positive SST correlation, whereas in the summer it does not (Fig. 9a, b). The wintertime tripole pattern features a negative SST correlation in the regions of the Caribbean to the western tropical North Atlantic and of the North Atlantic, and a positive correlation from the Gulf of Mexico to the southeastern coast of the United States (Fig. 9a). This difference may reflect that the NAO links the Icelandic low and the NASH and that the NAO is strong in the winter. The CLLJ is largely affected by variability of the NASH. When the NASH is strong and extends westward, the easterly CLLJ becomes strong. The strong easterly CLLJ results in a decrease in SST through evaporation and/or ocean dynamics in the CLLJ region and in the western tropical North Atlantic. In the region from the Gulf of Mexico to the southeastern coast of the United States, the CLLJ-related SLP pattern (Fig. 8a) suggests a westerly wind anomaly that can warm SST (Fig. 9a). The negative SST correlation at the high latitudes of around 45°N–50°N may involve the NAO, which represents a meridional SLP seesaw between the Azores high (or the NASH) and the Icelandic low (Hurrel 1995). When the NASH is strengthened (the easterly CLLJ is strong), the Icelandic low is strengthened (Fig. 8a). This corresponds to strong westerly wind in the high latitudes, and thus decreasing SST there (Fig. 9a).
4.4 Relation to anomalous meridional flows to the United States
The mechanism of the meridional wind correlations with the CLLJ can be explained in terms of atmospheric response to the oceanic SST forcing. A cold SST in the Caribbean region strengthens the NASH that in turn increases the easterly CLLJ. The pressure response in Fig. 8 shows a high SLP located north and northeast of the Caribbean region. This high SLP pattern produces southerly (northerly) wind anomalies at its western (eastern) side, thus resulting in a positive (negative) meridional wind correlation in the Gulf of Mexico and the central/eastern United States (the subtropical North Atlantic). The negative (positive) meridional wind correlation in the North Pacific in the winter (summer) in Fig. 11 represents northerly (southerly) wind anomalies that are associated with the wintertime Aleutian low (the summertime North Pacific subtropical high).
4.5 Relation to rainfall anomalies
In the summer, the CLLJ index is positively correlated with rainfall anomalies in the eastern North Pacific in a longitudinal band along 9°N (Fig. 12b). This positive correlation may also be associated with the CLLJ’s moisture transport. As the CLLJ passes the Caribbean Sea, one of the CLLJ’s branches is to continue westward across Central America into the eastern North Pacific. It is possible that a strong CLLJ enhances the moisture convergence in the eastern North Pacific and thus increases rainfall there. Figure 12 also shows that the eastern subtropical North Pacific displays significant rainfall correlations during both the winter and summer. This is consistent with the distributions of SLP anomalies and 925 hPa meridional wind anomalies in Figs. 8 and 11. In the winter, the North Pacific is under the influence of the Aleutian low. The positive SLP anomalies in the North Pacific (Fig. 8a) means a weakening of the Aleutian low which is associated with northerly wind anomalies (Fig. 11a). The northerly wind anomalies are not favorable for rainfall, resulting in negative rainfall correlation in the eastern subtropical North Pacific and the west coast of the United States (Fig. 12a). In the summer, the North Pacific is controlled by the North Pacific subtropical high. The negative SLP anomalies in the subtropical North Pacific (Fig. 8b) indicate a weakening of the subtropical high which is associated with southerly wind anomalies (Fig. 11b). The southerly wind anomalies carry moisture to the eastern subtropical North Pacific for more rainfall there (Fig. 12b). However, how the CLLJ is related to the Aleutian low and the North Pacific subtropical high is unknown.
The easterly CLLJ is observed to show a semi-annual variation, with two maxima in the summer and winter and two minima in the fall and spring.
The summertime strong easterly CLLJ is associated with a maximum of SLP, the MSD of rainfall, and a minimum of tropical cyclones in the Caribbean region. A possible mechanism for this relationship may be that the easterly CLLJ increases the moisture flux divergence in the Caribbean and thus suppresses the convection, decreasing rainfall, and suppressing the formation of tropical cyclones.
The semi-annual strengthening of the easterly CLLJ results from the semi-annual variation of the meridional SST and SLP gradients. A positive ocean–atmosphere feedback may be operating for maintaining the easterly CLLJ. A meridional SST gradient in the Caribbean induces a meridional SLP gradient that produces the easterly CLLJ. The easterly CLLJ in turn results in negative and positive wind stress curls to the north and south of the CLLJ core, respectively. The negative wind curl warms the northern Caribbean and the positive curl cools the southern Caribbean through oceanic Ekman dynamics, thus resulting in a further increase of the meridional SST gradient.
The CLLJ index of zonal wind anomalies at 925 hPa shows peaks of variability on timescales of 1.25, 2.3, and 10.2 years.
Two maxima of the standard deviation of CLLJ anomalies are found around September and February. The largest variance near September coincides with the busy and active months of Atlantic hurricanes, suggesting that the large variation of the easterly CLLJ changes the vertical wind shear between the lower and upper troposphere that then affects hurricane activity.
The CLLJ anomalies are inversely related to the Caribbean SST anomalies. Based on Gill’s theory, warm (cold) SST anomalies in the Caribbean are associated with low (high) SLP anomalies that weaken (strengthen) the easterly CLLJ.
The CLLJ varies in phase with the NAO. This reflects the fact that both the CLLJ and NAO are related to the NASH. A strong (weak) NASH is associated with strengthening (weakening) of the easterly CLLJ and also corresponds to the positive (negative) phase of the NAO.
The CLLJ has an opposite relationship with ENSO in the winter and summer. During the winter a weak (strong) easterly CLLJ corresponds to warm (cold) SST anomalies in the tropical Pacific, whereas during the summer a strong (weak) easterly CLLJ is associated with warm (cold) SST anomalies in the tropical Pacific. This is because ENSO’s teleconnections are different: ENSO induces negative and positive SLP anomalies in the subtropical North Atlantic during in the winter and summer, respectively.
The CLLJ is positively correlated with the surface meridional wind anomalies over the Gulf of Mexico and the central/eastern United States or the GPLLJ index. When the CLLJ is anomalously strong (weak), the meridional wind anomalies in the Gulf of Mexico and the central/eastern United States are southerly (northerly).
A strong (weak) easterly CLLJ is associated with southerly (northerly) wind anomalies to the United State that transport more (less) moisture for rainfall over the United States. At the same time, if more (less) moisture is exported from the ocean to the United States, it seems plausible that less (more) moisture would be available for local rainfall over the ocean region. Thus, the CLLJ and the GPLLJ are associated with an opposite rainfall pattern (i.e., a dipole rainfall pattern) in the tropical North Atlantic Ocean and Atlantic warm pool region versus the central United States.
I thank Mr Jay Harris for downloading the data used in this study. Dr Sang-ki Lee assists with Fig. 1 and calculates the significance test of the semi-annual feature of the CLLJ. Comments by Dr German Poveda, an anonymous reviewer, and the Editor (Dr Edwin Schneider) are appreciated. This work was supported by a grant from National Oceanic and Atmospheric Administration (NOAA) Climate Program Office and by the base funding of NOAA Atlantic Oceanographic and Meteorological Laboratory. The findings and conclusions in this report are those of the author(s) and do not necessarily represent the views of the funding agency.
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