EXTRATROPICAL CYCLONES 
dv,/dt being negative in the lower portion of the cold 
wedge. The sea-level map for November 9, 1500Z (Fig. 
10) shows a positive dv,/dx along the whole cold front 
(the geostrophic wind component parallel to the front 
in the cold air is increasingly negative as we pass 
towards the negative z-direction), but the actual wind 
component v,, parallel to the front, is about zero in 
the forward part of the cold wedge; dv,/dt also is 
small. Hence it follows that the isentropic downgliding 
v, Should be insignificant except where 20, — dv,/0n 
is zero or negative. The profile in Fig. 18 shows an al- 
most perfect parallelism of v,-isovels and dry-isentropes 
in the cold wedge, so that 0v,/dn = 0. Therefore, no 
dynamic instability occurs inside the cold air, not even 
in the upper part where the frontal zone is quite steep. 
The only place for dynamic instability to occur inside 
the cold air is right at the quasi-vertical part of the 
frontal surface near the ground, where a real temper- 
ature discontinuity exists. However, the air volume in- 
volved is too small to appear on a profile of the scale 
used in Fig. 13. The release of the dynamic instability 
at the cold front is responsible for maintaining the 
downdraft, which is always observed in a strip of a 
few kilometers’ width following the frontal passage. 
This cold-air downdraft is instrumental in giving the 
cold front a greater speed of displacement than would 
have been indicated by the geostrophic wind determined 
from the sea-level pressure distribution. In Fig. 10 
the computed twelve-hour geostrophic displacement of 
the cold front is represented by a short arrow of 80-km 
length, while at the same place the preceding twelve- 
hour displacement was 210 km and the subsequent one 
460 km. The geostrophic wind component normal to 
the front computed from the 850-mb map is large 
enough to account for the frontal displacement, and 
we must assume that air from this level enters the 
frontal downdraft and carries westerly momentum to 
the surface layer. In the layers above 850 mb the cold- 
air current is cyclonically curved and hence subgeo- 
strophic. With increasing height, the wind in the frontal 
zone also becomes more and more parallel to the frontal 
boundary, so that the frontal displacement is less there 
than at the ground. 
The prefrontal air in the profile in Fig. 13 has a 
lapse rate slightly less than the saturation adiabatic, 
but with the existing horizontal temperature gradient 
a saturation-adiabatic ascent is possible at the rather 
steep angle of 169, as shown by the sample saturation 
isentropes of 289° and 292°. The measured values of 
20, — dv,/dn show a close approach to dynamic in- 
difference and even some dynamic instability. A satu- 
ration-isentropic upgliding is in order at 850 mb, as 
can be seen from the convergence of contours (vz dv,/dx 
> 0) in the warm current intersected by the profile 
(Fig. 10). The same is true for the 700-mb map (not 
reproduced) but not for the 500-mb and 300-mb maps 
(Fig. 9). The frontal upgliding should therefore be con- 
fined to the layers under 500 mb. This is also verified 
by the Little Rock sounding, which goes up through the 
cold-front rain but shows a 35 per cent relative humidity 
at 500 mb. Thunderheads growing up from the cloud 
595 
mass of the cold front would, of course, go well beyond 
500 mb, but such phenomena of “‘vertical instability” 
were not reported in the case under consideration. 
Intermittent, light prefrontal ram, which was reported 
as far as 800 km ahead of the cold front, can be ac- 
counted for quite well by the saturation-isentropic 
upgliding. In the warm season such upgliding in the 
tropical air current may be sufficient to trigger thunder- 
storm formation, which in turn may develop prefrontal 
squall lines. 
Structure of the Maturing Frontal Cyclone. Figure 
14 presents the sea-level, 700-mb, 500-mb, and 300-mb 
maps for November 10, 0300Z, depicting the structure 
of the maturing frontal cyclone. The amplitude of the 
frontal wave on the surface map has now increased con- 
siderably, and the cold air from the rear begins to en- 
circle the cyclone center. It is clearly seen how this 
occlusion process has had its inception only at the 
ground, while the isotherm patterns of the upper maps 
still indicate an open wave of small amplitude. This 
shows that while the wave travels along the front the 
frontal slope increases to a maximum at the wave apex. 
At that point the smooth wave “breaks” and a rela- 
tively shallow cold outbreak fans out along the ground. 
The mechanism of this breaking of the smooth wave 
probably lies in the dynamic instability of the lower 
part of the frontal zone, described in its stage of in- 
ception in Fig. 12 and continuing in the form of the 
frontal downdraft in Fig. 13. During the process, the 
cold front part near the cyclone center undergoes fron- 
tolysis through cold-air downgliding. This is always 
noticeable in the surface-map analysis, and, on Novem- 
ber 10, 0300Z, it also shows up in a weakening of the 
frontal temperature gradient on the 700-mb map. 
Farther south, where the cold front passes through the 
frontogenetical field of deformation, the frontal tem- 
perature gradient remains rather strong. The strongest 
frontal temperature gradient, however, is found at the 
500-mb level where the breaking of the frontal wave 
has not yet started. The 300-mb map also shows fairly 
strong temperature gradients across the pressure trough, 
which may justify the use of the term “front” even at 
that level. Particularly striking is the crowding of iso- 
therms in the southwestern corner of the map, probably 
an. effect of frontogenesis in a 300-mb col in the un- 
mapped area west of Mexico. The 300-mb map under 
consideration is just tangent to a tropopause depression 
over the cold tongue of tropospheric air in the west. 
The map intersects the tropopause along the zone of 
lowest temperature across Labrador and northern 
Ontario. The temperature maximum east of the Hudson 
Bay cyclone is stratospheric. The cold air to its north 
is tropospheric and has been brought from lower lati- 
tudes as part of the warm sector shown in Fig. 9. On 
the 500-mb map, which is entirely tropospheric, the 
Hudson Bay center has a cold core with warmer sur- 
roundings both to the north and to the south. 
The pressure minimum of the frontal wave at the 
Great Lakes continues up to 700 mb with some west- 
ward tilt, but at that level it has already shrunk so as 
to be a minor feature in the pressure field compared to 
