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Volcanoes, the stratosphere, and climate
Journal of Volcanology and Geothermal Research, 28 (1986) 247--255 Elsevier Science Publishers B.V., Amsterdam -- Printed in The Netherlands
247
VOLCANOES, THE STRATOSPHERE, AND CLIMATE
BRUCE M. JAKOSKY Laboratory for Atmospheric and Space Physics, University of Colorado, Boulder, CO 80309, USA (Received July 5, 1985; revised and accepted February 2, 1986)
ABSTRACT Jakosky, B.M., 1986. Volcanoes, the stratosphere, and climate. J. Volcanol. Geotherm. Res., 28: 247--255. A list of volcanic eruption plumes observed to ascend into or near the stratosphere since 1883 shows that the volcanoes divide readily into two groups, one at low and one at higher latitudes. A model for the rise of a buoyant volcanic plume rise as applied to volcanic eruptions is corrected for realistic temperature profiles and for the finite vertical extent of the resultant debris clouds. The utility of the model can be questioned, however, owing to the highly uncertain and variable nature of the efficiency of use of heat energy of buoyant rise. The observed correlation of stratospheric plumes with climatic effects indicates that those plumes nearer the equator have the largest impact on surface temperatures. Analysis of the observations also suggests that injection of debris into the stratosphere is more important in determining the effect on climate than either the total volcanic explosivity of the eruption or the actual height reached within the stratosphere.
INTRODUCTION The injection o f volcanic debris into t h e s t r a t o s p h e r e can act as a forcing f u n c t i o n f o r climate o w i n g to the ability o f the aerosol to remain in the s t r a t o s p h e r e for years and to affect the energy balance o f t h e a t m o s p h e r e and the surface (e.g., L a m b , 1 9 7 0 ; Pollack et al., 1 9 7 6 ; T o o n and Pollack, 1982). In o r d e r to u n d e r s t a n d the relationship b e t w e e n volcanic e r u p t i o n s and possible effects on climate, several factors need to be u n d e r s t o o d : the global d i s t r i b u t i o n o f e r u p t i o n s reaching the s t r a t o s p h e r e and their t e m p o r a l behavior; t h e altitudes r e a c h e d by volcanic plumes, and their d e p e n d e n c e on the specific c o n d i t i o n s o f the e r u p t i o n (especially the e f f i c i e n c y o f utilization o f t h e t h e r m a l e n e r g y c o n t a i n e d in t h e solid ejecta); the realtionship b e t w e e n p l u m e c o m p o s i t i o n and l o c a t i o n and the resulting stratospheric aerosol c o n t e n t and l o c a t i o n ; and the effects o f the aerosol on the climate. This p a p e r deals with the r e c o r d o f observed injections o f debris into the s t r a t o s p h e r e during the last c e n t u r y , with the physics o f the injection process, and with the empirical relationship b e t w e e n volcanic e r u p t i o n s and climate. In the n e x t section, a list o f volcanic e r u p t i o n s ascending into or very near to the 0377-0273/86/$03.50
© 1986 Elsevier Science Publishers B.V.
248 TABLE 1 V o l c a n i c e r u p t i o n s o b s e r v e d to r e a c h t h e s t r a t o s p h e r e since 1 8 8 3 Symbol
Volcano
Date
Height (km)
A * B
Krakatau Etna
27 Aug 1 8 8 3 21 May 1 8 8 6
27 14
6 38
Lamb (1970) Lamb (1970)
C * D *
Soufri~re S a n t a Maria
17 May 1 9 0 2 24 Oct 1902
16 27--29
13 15
Wilson et al. ( 1 9 7 8 ) Rose ( 1 9 7 2 )
E
Quizapu
10 A p r 1932
14
45
Wilcox ( 1 9 5 9 )
F
Hekla
29 Mar 1947
26--27
64
Wilcox ( 1 9 5 9 )
G H I
Trident Spurr Bezymianny
15 F e b 1 9 5 2 9 Jul 1 9 5 3 30 Mar 1 9 5 6
9 22 36--45
58 61 56
Wilcox ( 1 9 5 9 ) Wilcox ( 1 9 5 9 ) Gorshkov (1959)
J K * L M N * O P Q *
Tokachidake Agung Trident Surtsey Taal Redoubt D e c e p t i o n Is. Fernandina
29 17 1 14 28 9 5 11
12 22 15 14.5 15--20 12--16 10 22--24
43 8 58 63 14 60 63 '/2
Walker ( 1 9 8 1 ) Dyer a n d Hicks ( 1 9 6 8 ) Decker (1967) Cronin (1971) M o o r e et al. ( 1 9 6 6 ) Decker (1967) Lamb (1970) Simkin and Howard (1970)
R
Hekla
16
64
S* T U *
Fuego 14 O c t 1974 St. A u g u s t i n e 23 J a n 1 9 7 6 SoufriSre 22 Apr 1 9 7 9
22 11 17--19
14,5 59 13
Thorarinsson and Sigvaldason ( 1 9 7 2 ) McCormicketal.(1978) Meinel et al. ( 1 9 7 6 ) Sparks a n d Wilson ( 1 9 8 2 )
V
St. Helens
18 May 1 9 8 0
24
46
W
St. Helens
7 Aug 1 9 8 0
13
46
X
St. Helens
18 O c t 1 9 8 0
14
46
Y Z * a *
Alaid Pagan "mystery cloud" E1 C h i c h o n
28 A p r 1981 15 May 1981 Jan 1982
15 13.5 17
/3 *
J a n 1962 Mar 1 9 6 3 Apr 1963 Nov 1963 Sep 1 9 6 5 Feb 1966 Jul 1967 J u n 1968
5 May 1 9 7 0
4 Apr 1982
28
*Marks t h o s e v o l c a n o e s w i t h i n 20 ° o f t h e e q u a t o r .
Lat
51 18 ~- 20 17
Reference
Christiansen and Peterson (1981) C h r i s t i a n s e n and Peterson (1981) Christiansen and Peterson (1981) Hofmann and Rosen (1981) EOS (1981) M.P. M c C o r m i c k (pets. c o m m u n . , 1 9 8 3 ) T h o m a s et al. ( 1 9 8 3 )
249 stratosphere since 1883 is presented. In the third section, a model of volcanic plume rise through the atmosphere developed by Morton et al. (1956) and by Wilson et al. {1978) is considered. Finally, the relationship between volcanic eruptions reaching the stratosphere and the effects on climate is discussed. OBSERVATIONS OF VOLCANIC PLUMES Table 1 lists twenty-eight volcanic plumes that were observed to have reached into or very near to the stratosphere. The list is not exhaustive, and includes only those eruptions for which there existed a direct observation of the volcanic plume itself or of the resulting stratospheric aerosol cloud. Table 1 was compiled from previously published analyses, and includes the volcanoes listed by Cronin (1971), Wilson et al. (1978), and others. This list differs from that compiled by Newhall and Self {1982), which was based on a Volcanic Explosivity Index (VEI) and which used several factors to gauge the magnitude of an eruption, including ejecta volume, rate of eruption, plume height, and destructive power. It also differs from that compiled by Lamb (1970), which was based predominantly on the resultant temperature effects of stratospheric aerosol clouds. By limiting Table 1 to those eruptions which were observed to have reached the stratosphere, the effects of the eruptions on climate may be better discerned. Being based purely on observed plumes, some relevant eruptions may not have been included in Table 1. An example of a significant eruption which has not been included is the 1912 eruption of Katmai, which had a notable effect on temperatures but for which there was no direct observation (Lamb, 1970). Newhall and Self (1982) list fifty-one "cataclysmic" or greater 50
[
i
T I
,-,4O
1
E }.r30 ~D
A D/~
F
LU LU 20
~-rr----~-~ ~ . !. R Q..
"J"~.qQT -P
10
0
30
60
90
LATITUDE
Fig. 1. The plume height reached and latitude for each volcanic eruption listed in Table 1.
The seasonal extremes of Table 1.
t h e a l t i t u d e to t h e t r o p o p a u s e
are
marked. Symbols are from
250 (VEI ~ 4) eruptions during the same time period as encompassed by Table 1; however, the VEI is based on factors in addition to plume height, such that all of these eruptions may not have reached the stratosphere. Increased observations from the ground and from spacecraft have recently enhanced the ability to detect the resultant stratospheric aerosol cloud; some eruptions have been described which might otherwise have been missed (e.g., the so-called "mystery cloud" of 1982). In Table 1, the altitude of the 1963 Agung eruption refers to the resultant aerosol cloud (22 km, Dyer and Hicks, 1968) rather than to the plume height observed visually ( 1 0 k m , Settle, 1978). Figure 1 shows the plume height and latitude of the eruptions listed in Table 1, along w i & the seasonal extremes of the altitude to the tropopause. As noted by Cronin (1971), most of the eruptions fall either within 20 ° of the equator or between 40 and 70 ° from the equator. Implications of this distribution will be discussed in the last section. MODEL OF PLUME RISE The ability of a volcanic eruption to inject debris into the stratosphere depends on the rate of injection of heat, as the plume rises due to its buoyancy with respect to the surrounding ambient atmosphere (Morton et al., 1956; Briggs, 1969; Wilson et al., 1978). The classic model of plume rise developed by Morton et al. (1956) has been applied to volcanic eruptions by Briggs (1969) and by Wilson et al. (1978). The plume rise height is given for a steady plume by: H = 31(1 + t2)-3/8(~ 1/4
(1)
where H is in meters, Q is the source rate of thermal energy production in kilowatts, and n is the ratio of the actual vertical temperature gradient to the adiabatic lapse rate, taken to be constant with altitude. Wilson et al. (1978) relate Q to the eruption-specific conditions at the vent via: = ~u~r2s(O --Oa)F
(2)
where ~, u, s, and 0 are the bulk density, velocity, specific heat, and temperature of the erupting material, 0a is the temperature to which the eruption products ultimately cool, r is the vent radius, and F is an efficiency factor for the energy usage. For the cases where O has been inferred or measured, the plume height appears to follow the Q1/4 behavior of eqn. (1) quite well (Wilson et al. 1978; Settle, 1978). This agreement may be fortuitous, however, for the following reasons: (1) Equation (1) was derived assuming a constant lapse rate appropriate for the troposphere (see Briggs, 1969; and Morton et al., 1956). The increasing temperatures within the stratosphere will result in plume rise to a lower altitude due to the decrease in thermal c~,~rast between the plume and the ambient atmosphere. (2) Equation (1) was derived as being descriptive of the rise of the bulk
251
of the plume material, and should therefore be applied to the mass-average height to which the plume rises, yet it has been applied to the visible top of the plume by Settle (1978) and Wilson et al. (1978). These t w o affects, discussed below in detail, can result in a greater than 10 km change in the predicted plume rise height. A simple correction is made here to account for the non-linear temperature profile. For each altitude, the value of Q in eqn. (1) that is required was calculated using the average atmospheric lapse rate from the surface up to that altitude. Ambient air at different altitudes will mix with the plume in varying proportions, with air from the lower atmosphere entrained preferentially owing to the larger surface area to volume ratio of the plume at low altitudes (L. Wilson, pers. commun., 1985); this calculation is appropriate, however, because the height to which the plume rises depends on the contrast between the plume density and the density of the ambient air at each altitude. Figure 2 compares the observed volcanic plume heights and eruption rates with eqn. (1) (from Wilson et al., 1978, with F = 1) and with the plume heights predicted from this modified version of eqn. (1). Examination of Fig. 2 suggests that the plume height should vary as O0.1s for heights above the tropopause, rather than varying as O0.2s. At high altitudes, this effect can lead to a 15-km difference in predicted plume height. 50
v /
4O
/
/
//
/
/
j
/
/
/
/*
/
Bezymianny-I-// /"
/
/ t
/ / / v
j
z
/ /
z
/ /
~
/* f /
30 S a l t aj Marina +
//-
I.-I-
/Y
LU "I-
~/
Hekla +
./
/
/ "e~ / ~ o ~, ..-/.xX ./
-"
-t- Agung / / /
uJ 20
/
~¢//
/ / /
+Soufriere
,/
Hekla +/. / / / / J /
lO¢"
loa
J
I
104
10 s
VOLUME ERUPTION RATE, ma/s
Fig. 2. Observations o f p l u m e height and eruption rate (from Settle, 1978; and Wilson et al., 1 9 7 8 ) compared w i t h the m o d e l by Wilson et al. (heavy line). The curve labeled "Modified" includes the non-constant atmospheric lapse rate, as discussed in the text. The light-dashed curves have been increased b y 50% to a c c o u n t for the finite vertical e x t e n t o f the clouds.
252
The second effect is more difficult to predict owing to the large variability of debris rise even within an individual plume. For instance, the eruption of Bezymianny in 1956 resulted in a plume that was distributed in altitude between about 8 and 36kin, with only one small part going as high as 45 km (Gorshkov, 1959). To estimate the role of the finite vertical extent of the plume, we follow Briggs (1969) and assume that, to first order, the total vertical extent of the emplaced cloud is roughly equal to the average height reached. This effect is shown in Fig. 2 by increasing the plume height reached by 50%. Notice that all of the eruptions except Agung now fall below the predicted curve, as would be expected for an efficiency factor less than 1.0. Even Agung will fall below the model if the plume height observed visually, about 10 kin, were used in place of the aerosol cloud height. The utility of using this model is limited, however, by the large uncertainties in the measurement of the volumetric rate of eruption as well as in the uncertainty of the efficiency factor. The efficiency factor F depends on a variety of factors, including the degree of entrainment of solid ejecta and the specific geometry of the eruption. For example, a plume with a high density of solid debris can fail to be supported by buoyancy alone and collapse under its own weight (Wilson et al., 1978). EFFECTS ON CLIMATE Cronin (1971) suggests that the high4atitude eruptions (Fig. 1) may be more important in affecting the stratospheric aerosol than the low-latitude eruptions owing to the lower altitude of the tropopause at high latitude and the consequent greater ease of injection. This effect can be seen in SAGE satellite observations of the aerosol from the Mt. St. Helens eruption of 1980 (Inn et al., 1982), which show that most of the aerosol in that case lies between about 10 and 15 km altitude; had the eruption been at the equator, this material probably would not have reached the stratosphere. It is important to note, however, that the stratospheric circulation may be so sluggish that much of the high-latitude material may never reach the equator; this appears to be the case for the Mount St. Helen's aerosol, which remained confined to high latitudes (see Inn et al., 1982). An additional factor in determining the effects of the stratospheric aerosol is the abundance of sulfur in the eruptive material. The major part of the aerosol is predominantly sulfuric acid, produced from nucleation of SO 2 gas and growth of aerosol droplets (e.g., Turco, 1982). An eruption like that of E1 Chichon in 1982, being rich in sulfur (e.g., Varekamp et al., 1984), will produce a larger aerosol cloud than a greater eruption with lessabundant sulfur. By way of comparison, the eruptions of Mt. St. Helens (1980) and E1 Chichon (1982) were comparable in magnitude, yet more stratospheric aerosol was produced by the latter, and it was much more significant climatically (see Pollack and Ackerman, 1983).
253
Comparison of the volcanic plumes reaching the stratosphere with the dust veil index of Lamb (1970) shows a remarkably good correlation (Fig. 3). Lamb's (1970) dust veil index-is predominantly a measure of the effect of volcanic eruptions on climate. It is interesting to notice that the largest increases in Lamb's (1970) index correspond to the eruption of volcanoes nearer to the equator. This is an important result in understanding the role of volcanoes in affecting climate, but may also be an artifact of where on the globe temperatures are measured. Also, there is a better correlation of climatic effects with the occurrence of volcanic plumes observed to reach the stratospheric than there is with the occurrence of large eruptions as indicated by the Volcanic Explosivity Index of Newhall and Self (1982.) This result is reasonable given the large variety of factors included in the VEI. It is apparent from Fig. 3 that volcanic eruptions do have an effect on climate, that the effect correlates with the location of the eruption but not necessarily with the height reached by the plume, and that it does not correlate with a gross measure of the explosivity of the eruption. Factors such as the efficiency of energy usage (which is tied closely to the mass loading of the plume and the geometry of the vent) and the sulfur c o n t e n t of the erupting fluid are not yet well understood; they certainly play a major role in providing the large variability shown in Figs. 2 and 3.
(c]
>
VV
I)
I
[~
T
I r)l
N
iii i i
vV
Iii
¢
N ~
I
iii
ilfl
V
i iii
itll
v
i
v (b)
r~rrw~
~ o ~, Z.j Z
Z
,','x u. ~
~o 600
VV
~
V
V
~V
V VV (a)
400
~ ~
200 0 188
1890
1900
1910
1920
1930
1940
1950
1960
1970
1980
1990
YEAR
Fig. 3. (a) The dust veil i n d e x o f L a m b (1970) since 1980, based o n o b s e r v e d aerosols and their effects o n surface temperatures. (b) The year ot~ the eruption o f those volcanoes f r o m Table 1. Those w i t h i n 20 ° o f the equator are m a r k e d b y filled triangles and have their n a m e s listed. (c) V o l c a n o e s from the list o f Newhall and Self (1982) w h i c h have V E I / > 5 (triangles) and V E I ----4 (lines).
254 REFERENCES Briggs, G.A., 1969. Plume rise. Critical Review Series, Rep. TID-25075, Atom. Energy Comm., Washington, D.C. Christiansen, R.L. and Peterson, D.W., 1981. Chronology of the 1980 eruptive activity. In: The 1980 Eruptions of Mount St. Helens, Washington. U.S. Geol. Surv., Prof. Pap., 1250: 17--30. Cronin, J.F., 1971. Recent volcanism and the stratosphere. Science, 172: 847--849. Decker, R.W., 1967. Investigations at active volcanoes. Trans. Am. Geophys. Union, 48 : 639--647. Dyer, A.J. and Hicks, B.B., 1968. Global spread of volcanic dust from the Bali eruption of 1963. Q. J. R. Meteorol. Soc., 94: 545--554. EOS, Trans. Am. Geophys. Union, 1981, 62: 510. Gorshkov, G.S., 1959. Gigantic eruptions of the volcano Bezymianny. Bull. Volcanol., 20: 77--109. Hofmann, D.J. and Rosen, J.M., 1981. Stratospheric aerosol and condensation nuclei enhancements following the eruption of Alaid in April 1981. Geophys. Res. Lett., 8: 1231--1234. Inn, E.C.Y., Farlow, M.H., Russel, P.B., McCormick, M.P. and Chu, W.P., 1982. Observations. In: R.C. Whitten (Editor), The Stratospheric Aerosol Layer. Springer-Verlag, New York, N.Y., pp. 15--68. Lamb, H.H., 1970. Volcanic dust in the atmosphere; with a chronology and assessment of its meteorological significance. Philos. Trans. R. Soc. London, Set. A, 266: 425-553. McCormick, M.P., Swissler, T.J., Chu, W.P. and Fuller, W.H., Jr., 1978. Post-volcanic stratospheric aerosol decay as measured by lidar. J. Atmos. Sci., 35: 1296--1303. Meinel, A.B., Meinel, M.P. and Shaw, G.E., 1976. Trajectory of the Mt. St. Augustine 1976 eruption ash cloud. Science, 193: 420--422. Moore, J.G., Nakamura, K. and Alcarez, A., 1966. The 1965 eruption of Taal volcano. Science, 151: 955--960. Morton, B.R., Taylor, G. and Turner, J.S., 1956. Turbulent gravitational convection from maintained and instantaneous sources. Proc. R. Soc., Set. A., 234: 1--23. Newhall, C.G. and Self, S., 1982. The volcanic explosivity index (V.E.I.): An estimate of explosive magnitude for historical volcanism. J. Geophys. Res., 87: 1231--1238. Pollack, J.B. and Ackerman, T.P., 1983. Possible effects of the'E1 Chichon volcanic cloud on the radiation budget of the northern tropics. Geophys. Res. Lett., 10: 1057--1060. Pollack, J.B., Toon, O.B., Sagan, C., Summers, A., Baldwin, B. and Van Camp, W., 1976. Volcanic explosions and climatic change: A theoretical assessment. J. Geophys. Res., 81: 1071--1083. Rose, W.I., Jr., 1972. Notes on the 1902 eruption of Santa Maria volcano, Guatemala. Bull. Volcanol., 36: 29--45. Settle, M., 1978. Volcanic eruption clouds and the thermal power output of explosive eruptions. J. Volcanol. Geotherm. Res., 3: 309--324. Simkin, T. and Howard, K.A., 1970. Caldera collapse in the Galapagos Islands, 1968. Science, 169: 429--437. Sparks, R.S.J. and Wilson, L., 1982. Explosive volcanic eruptions -- V. Observations of plume dynamics during the 1979 Soufri~re eruption, St. Vincent. Geophys. J.R. Astron. Soc, 69: 551--570. Thomas, G.E., Jakosky, B.M., West, R.A. and Sanders, R.W., 1983. Satellite limbscanning thermal infrared observations of the E1 Chichon stratospheric aerosol: First results. Geophys. Res. Lett., 10: 997--1000. Thorarinsson, S. and Sigvaldason, G.E., 1972. The Hekla eruption of 1970. Bull. Volcanol., 36: 269--288.
255 Toon, O.B. and Pollack, J.B., 1982. Stratospheric aerosols and climate. In: R.C. Whitten (Editor), The Stratospheric Aerosol Layer. Springer-Verlag, New York, pp. 121--147. Turco, R.P., 1982. Models of stratospheric aerosols and dust. In: R.C. Whitten (Editor), The Stratospheric Aerosol Layer. Springer-Verlag, New York, pp. 93--119. Varekamp, J.C., Luhr, J.F. and Prestegaard, K.L., 1984. The 1982 eruptions of E1Chichon volcano (Chiapas, Mexico): Character of the eruption, ash-fall deposits, and gas phase. J. Volcanol. Geotherm. Res., 23: 39--68. Walker, G.P.L., 1981. Generation and dispersal of fine ash and dust by volcanic eruptions. J. Volcanol. Geothermal Res., 11 : 81--92. Wilcox, R.E., 1959. Some effects of recent volcanic ash falls with especial reference to Alaska. U.S. Geol. Surv., Bull., 1028-N: 409--476. Wilson, L., Sparks, R.S.J., Huang, T.C. and Watkins, N.D., 1978. The control of volcanic column heights by eruption energetics and dynamics. J. Geophys. Res., 83: 1829-1836.
247
VOLCANOES, THE STRATOSPHERE, AND CLIMATE
BRUCE M. JAKOSKY Laboratory for Atmospheric and Space Physics, University of Colorado, Boulder, CO 80309, USA (Received July 5, 1985; revised and accepted February 2, 1986)
ABSTRACT Jakosky, B.M., 1986. Volcanoes, the stratosphere, and climate. J. Volcanol. Geotherm. Res., 28: 247--255. A list of volcanic eruption plumes observed to ascend into or near the stratosphere since 1883 shows that the volcanoes divide readily into two groups, one at low and one at higher latitudes. A model for the rise of a buoyant volcanic plume rise as applied to volcanic eruptions is corrected for realistic temperature profiles and for the finite vertical extent of the resultant debris clouds. The utility of the model can be questioned, however, owing to the highly uncertain and variable nature of the efficiency of use of heat energy of buoyant rise. The observed correlation of stratospheric plumes with climatic effects indicates that those plumes nearer the equator have the largest impact on surface temperatures. Analysis of the observations also suggests that injection of debris into the stratosphere is more important in determining the effect on climate than either the total volcanic explosivity of the eruption or the actual height reached within the stratosphere.
INTRODUCTION The injection o f volcanic debris into t h e s t r a t o s p h e r e can act as a forcing f u n c t i o n f o r climate o w i n g to the ability o f the aerosol to remain in the s t r a t o s p h e r e for years and to affect the energy balance o f t h e a t m o s p h e r e and the surface (e.g., L a m b , 1 9 7 0 ; Pollack et al., 1 9 7 6 ; T o o n and Pollack, 1982). In o r d e r to u n d e r s t a n d the relationship b e t w e e n volcanic e r u p t i o n s and possible effects on climate, several factors need to be u n d e r s t o o d : the global d i s t r i b u t i o n o f e r u p t i o n s reaching the s t r a t o s p h e r e and their t e m p o r a l behavior; t h e altitudes r e a c h e d by volcanic plumes, and their d e p e n d e n c e on the specific c o n d i t i o n s o f the e r u p t i o n (especially the e f f i c i e n c y o f utilization o f t h e t h e r m a l e n e r g y c o n t a i n e d in t h e solid ejecta); the realtionship b e t w e e n p l u m e c o m p o s i t i o n and l o c a t i o n and the resulting stratospheric aerosol c o n t e n t and l o c a t i o n ; and the effects o f the aerosol on the climate. This p a p e r deals with the r e c o r d o f observed injections o f debris into the s t r a t o s p h e r e during the last c e n t u r y , with the physics o f the injection process, and with the empirical relationship b e t w e e n volcanic e r u p t i o n s and climate. In the n e x t section, a list o f volcanic e r u p t i o n s ascending into or very near to the 0377-0273/86/$03.50
© 1986 Elsevier Science Publishers B.V.
248 TABLE 1 V o l c a n i c e r u p t i o n s o b s e r v e d to r e a c h t h e s t r a t o s p h e r e since 1 8 8 3 Symbol
Volcano
Date
Height (km)
A * B
Krakatau Etna
27 Aug 1 8 8 3 21 May 1 8 8 6
27 14
6 38
Lamb (1970) Lamb (1970)
C * D *
Soufri~re S a n t a Maria
17 May 1 9 0 2 24 Oct 1902
16 27--29
13 15
Wilson et al. ( 1 9 7 8 ) Rose ( 1 9 7 2 )
E
Quizapu
10 A p r 1932
14
45
Wilcox ( 1 9 5 9 )
F
Hekla
29 Mar 1947
26--27
64
Wilcox ( 1 9 5 9 )
G H I
Trident Spurr Bezymianny
15 F e b 1 9 5 2 9 Jul 1 9 5 3 30 Mar 1 9 5 6
9 22 36--45
58 61 56
Wilcox ( 1 9 5 9 ) Wilcox ( 1 9 5 9 ) Gorshkov (1959)
J K * L M N * O P Q *
Tokachidake Agung Trident Surtsey Taal Redoubt D e c e p t i o n Is. Fernandina
29 17 1 14 28 9 5 11
12 22 15 14.5 15--20 12--16 10 22--24
43 8 58 63 14 60 63 '/2
Walker ( 1 9 8 1 ) Dyer a n d Hicks ( 1 9 6 8 ) Decker (1967) Cronin (1971) M o o r e et al. ( 1 9 6 6 ) Decker (1967) Lamb (1970) Simkin and Howard (1970)
R
Hekla
16
64
S* T U *
Fuego 14 O c t 1974 St. A u g u s t i n e 23 J a n 1 9 7 6 SoufriSre 22 Apr 1 9 7 9
22 11 17--19
14,5 59 13
Thorarinsson and Sigvaldason ( 1 9 7 2 ) McCormicketal.(1978) Meinel et al. ( 1 9 7 6 ) Sparks a n d Wilson ( 1 9 8 2 )
V
St. Helens
18 May 1 9 8 0
24
46
W
St. Helens
7 Aug 1 9 8 0
13
46
X
St. Helens
18 O c t 1 9 8 0
14
46
Y Z * a *
Alaid Pagan "mystery cloud" E1 C h i c h o n
28 A p r 1981 15 May 1981 Jan 1982
15 13.5 17
/3 *
J a n 1962 Mar 1 9 6 3 Apr 1963 Nov 1963 Sep 1 9 6 5 Feb 1966 Jul 1967 J u n 1968
5 May 1 9 7 0
4 Apr 1982
28
*Marks t h o s e v o l c a n o e s w i t h i n 20 ° o f t h e e q u a t o r .
Lat
51 18 ~- 20 17
Reference
Christiansen and Peterson (1981) C h r i s t i a n s e n and Peterson (1981) Christiansen and Peterson (1981) Hofmann and Rosen (1981) EOS (1981) M.P. M c C o r m i c k (pets. c o m m u n . , 1 9 8 3 ) T h o m a s et al. ( 1 9 8 3 )
249 stratosphere since 1883 is presented. In the third section, a model of volcanic plume rise through the atmosphere developed by Morton et al. (1956) and by Wilson et al. {1978) is considered. Finally, the relationship between volcanic eruptions reaching the stratosphere and the effects on climate is discussed. OBSERVATIONS OF VOLCANIC PLUMES Table 1 lists twenty-eight volcanic plumes that were observed to have reached into or very near to the stratosphere. The list is not exhaustive, and includes only those eruptions for which there existed a direct observation of the volcanic plume itself or of the resulting stratospheric aerosol cloud. Table 1 was compiled from previously published analyses, and includes the volcanoes listed by Cronin (1971), Wilson et al. (1978), and others. This list differs from that compiled by Newhall and Self {1982), which was based on a Volcanic Explosivity Index (VEI) and which used several factors to gauge the magnitude of an eruption, including ejecta volume, rate of eruption, plume height, and destructive power. It also differs from that compiled by Lamb (1970), which was based predominantly on the resultant temperature effects of stratospheric aerosol clouds. By limiting Table 1 to those eruptions which were observed to have reached the stratosphere, the effects of the eruptions on climate may be better discerned. Being based purely on observed plumes, some relevant eruptions may not have been included in Table 1. An example of a significant eruption which has not been included is the 1912 eruption of Katmai, which had a notable effect on temperatures but for which there was no direct observation (Lamb, 1970). Newhall and Self (1982) list fifty-one "cataclysmic" or greater 50
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The seasonal extremes of Table 1.
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250 (VEI ~ 4) eruptions during the same time period as encompassed by Table 1; however, the VEI is based on factors in addition to plume height, such that all of these eruptions may not have reached the stratosphere. Increased observations from the ground and from spacecraft have recently enhanced the ability to detect the resultant stratospheric aerosol cloud; some eruptions have been described which might otherwise have been missed (e.g., the so-called "mystery cloud" of 1982). In Table 1, the altitude of the 1963 Agung eruption refers to the resultant aerosol cloud (22 km, Dyer and Hicks, 1968) rather than to the plume height observed visually ( 1 0 k m , Settle, 1978). Figure 1 shows the plume height and latitude of the eruptions listed in Table 1, along w i & the seasonal extremes of the altitude to the tropopause. As noted by Cronin (1971), most of the eruptions fall either within 20 ° of the equator or between 40 and 70 ° from the equator. Implications of this distribution will be discussed in the last section. MODEL OF PLUME RISE The ability of a volcanic eruption to inject debris into the stratosphere depends on the rate of injection of heat, as the plume rises due to its buoyancy with respect to the surrounding ambient atmosphere (Morton et al., 1956; Briggs, 1969; Wilson et al., 1978). The classic model of plume rise developed by Morton et al. (1956) has been applied to volcanic eruptions by Briggs (1969) and by Wilson et al. (1978). The plume rise height is given for a steady plume by: H = 31(1 + t2)-3/8(~ 1/4
(1)
where H is in meters, Q is the source rate of thermal energy production in kilowatts, and n is the ratio of the actual vertical temperature gradient to the adiabatic lapse rate, taken to be constant with altitude. Wilson et al. (1978) relate Q to the eruption-specific conditions at the vent via: = ~u~r2s(O --Oa)F
(2)
where ~, u, s, and 0 are the bulk density, velocity, specific heat, and temperature of the erupting material, 0a is the temperature to which the eruption products ultimately cool, r is the vent radius, and F is an efficiency factor for the energy usage. For the cases where O has been inferred or measured, the plume height appears to follow the Q1/4 behavior of eqn. (1) quite well (Wilson et al. 1978; Settle, 1978). This agreement may be fortuitous, however, for the following reasons: (1) Equation (1) was derived assuming a constant lapse rate appropriate for the troposphere (see Briggs, 1969; and Morton et al., 1956). The increasing temperatures within the stratosphere will result in plume rise to a lower altitude due to the decrease in thermal c~,~rast between the plume and the ambient atmosphere. (2) Equation (1) was derived as being descriptive of the rise of the bulk
251
of the plume material, and should therefore be applied to the mass-average height to which the plume rises, yet it has been applied to the visible top of the plume by Settle (1978) and Wilson et al. (1978). These t w o affects, discussed below in detail, can result in a greater than 10 km change in the predicted plume rise height. A simple correction is made here to account for the non-linear temperature profile. For each altitude, the value of Q in eqn. (1) that is required was calculated using the average atmospheric lapse rate from the surface up to that altitude. Ambient air at different altitudes will mix with the plume in varying proportions, with air from the lower atmosphere entrained preferentially owing to the larger surface area to volume ratio of the plume at low altitudes (L. Wilson, pers. commun., 1985); this calculation is appropriate, however, because the height to which the plume rises depends on the contrast between the plume density and the density of the ambient air at each altitude. Figure 2 compares the observed volcanic plume heights and eruption rates with eqn. (1) (from Wilson et al., 1978, with F = 1) and with the plume heights predicted from this modified version of eqn. (1). Examination of Fig. 2 suggests that the plume height should vary as O0.1s for heights above the tropopause, rather than varying as O0.2s. At high altitudes, this effect can lead to a 15-km difference in predicted plume height. 50
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Fig. 2. Observations o f p l u m e height and eruption rate (from Settle, 1978; and Wilson et al., 1 9 7 8 ) compared w i t h the m o d e l by Wilson et al. (heavy line). The curve labeled "Modified" includes the non-constant atmospheric lapse rate, as discussed in the text. The light-dashed curves have been increased b y 50% to a c c o u n t for the finite vertical e x t e n t o f the clouds.
252
The second effect is more difficult to predict owing to the large variability of debris rise even within an individual plume. For instance, the eruption of Bezymianny in 1956 resulted in a plume that was distributed in altitude between about 8 and 36kin, with only one small part going as high as 45 km (Gorshkov, 1959). To estimate the role of the finite vertical extent of the plume, we follow Briggs (1969) and assume that, to first order, the total vertical extent of the emplaced cloud is roughly equal to the average height reached. This effect is shown in Fig. 2 by increasing the plume height reached by 50%. Notice that all of the eruptions except Agung now fall below the predicted curve, as would be expected for an efficiency factor less than 1.0. Even Agung will fall below the model if the plume height observed visually, about 10 kin, were used in place of the aerosol cloud height. The utility of using this model is limited, however, by the large uncertainties in the measurement of the volumetric rate of eruption as well as in the uncertainty of the efficiency factor. The efficiency factor F depends on a variety of factors, including the degree of entrainment of solid ejecta and the specific geometry of the eruption. For example, a plume with a high density of solid debris can fail to be supported by buoyancy alone and collapse under its own weight (Wilson et al., 1978). EFFECTS ON CLIMATE Cronin (1971) suggests that the high4atitude eruptions (Fig. 1) may be more important in affecting the stratospheric aerosol than the low-latitude eruptions owing to the lower altitude of the tropopause at high latitude and the consequent greater ease of injection. This effect can be seen in SAGE satellite observations of the aerosol from the Mt. St. Helens eruption of 1980 (Inn et al., 1982), which show that most of the aerosol in that case lies between about 10 and 15 km altitude; had the eruption been at the equator, this material probably would not have reached the stratosphere. It is important to note, however, that the stratospheric circulation may be so sluggish that much of the high-latitude material may never reach the equator; this appears to be the case for the Mount St. Helen's aerosol, which remained confined to high latitudes (see Inn et al., 1982). An additional factor in determining the effects of the stratospheric aerosol is the abundance of sulfur in the eruptive material. The major part of the aerosol is predominantly sulfuric acid, produced from nucleation of SO 2 gas and growth of aerosol droplets (e.g., Turco, 1982). An eruption like that of E1 Chichon in 1982, being rich in sulfur (e.g., Varekamp et al., 1984), will produce a larger aerosol cloud than a greater eruption with lessabundant sulfur. By way of comparison, the eruptions of Mt. St. Helens (1980) and E1 Chichon (1982) were comparable in magnitude, yet more stratospheric aerosol was produced by the latter, and it was much more significant climatically (see Pollack and Ackerman, 1983).
253
Comparison of the volcanic plumes reaching the stratosphere with the dust veil index of Lamb (1970) shows a remarkably good correlation (Fig. 3). Lamb's (1970) dust veil index-is predominantly a measure of the effect of volcanic eruptions on climate. It is interesting to notice that the largest increases in Lamb's (1970) index correspond to the eruption of volcanoes nearer to the equator. This is an important result in understanding the role of volcanoes in affecting climate, but may also be an artifact of where on the globe temperatures are measured. Also, there is a better correlation of climatic effects with the occurrence of volcanic plumes observed to reach the stratospheric than there is with the occurrence of large eruptions as indicated by the Volcanic Explosivity Index of Newhall and Self (1982.) This result is reasonable given the large variety of factors included in the VEI. It is apparent from Fig. 3 that volcanic eruptions do have an effect on climate, that the effect correlates with the location of the eruption but not necessarily with the height reached by the plume, and that it does not correlate with a gross measure of the explosivity of the eruption. Factors such as the efficiency of energy usage (which is tied closely to the mass loading of the plume and the geometry of the vent) and the sulfur c o n t e n t of the erupting fluid are not yet well understood; they certainly play a major role in providing the large variability shown in Figs. 2 and 3.
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YEAR
Fig. 3. (a) The dust veil i n d e x o f L a m b (1970) since 1980, based o n o b s e r v e d aerosols and their effects o n surface temperatures. (b) The year ot~ the eruption o f those volcanoes f r o m Table 1. Those w i t h i n 20 ° o f the equator are m a r k e d b y filled triangles and have their n a m e s listed. (c) V o l c a n o e s from the list o f Newhall and Self (1982) w h i c h have V E I / > 5 (triangles) and V E I ----4 (lines).
254 REFERENCES Briggs, G.A., 1969. Plume rise. Critical Review Series, Rep. TID-25075, Atom. Energy Comm., Washington, D.C. Christiansen, R.L. and Peterson, D.W., 1981. Chronology of the 1980 eruptive activity. In: The 1980 Eruptions of Mount St. Helens, Washington. U.S. Geol. Surv., Prof. Pap., 1250: 17--30. Cronin, J.F., 1971. Recent volcanism and the stratosphere. Science, 172: 847--849. Decker, R.W., 1967. Investigations at active volcanoes. Trans. Am. Geophys. Union, 48 : 639--647. Dyer, A.J. and Hicks, B.B., 1968. Global spread of volcanic dust from the Bali eruption of 1963. Q. J. R. Meteorol. Soc., 94: 545--554. EOS, Trans. Am. Geophys. Union, 1981, 62: 510. Gorshkov, G.S., 1959. Gigantic eruptions of the volcano Bezymianny. Bull. Volcanol., 20: 77--109. Hofmann, D.J. and Rosen, J.M., 1981. Stratospheric aerosol and condensation nuclei enhancements following the eruption of Alaid in April 1981. Geophys. Res. Lett., 8: 1231--1234. Inn, E.C.Y., Farlow, M.H., Russel, P.B., McCormick, M.P. and Chu, W.P., 1982. Observations. In: R.C. Whitten (Editor), The Stratospheric Aerosol Layer. Springer-Verlag, New York, N.Y., pp. 15--68. Lamb, H.H., 1970. Volcanic dust in the atmosphere; with a chronology and assessment of its meteorological significance. Philos. Trans. R. Soc. London, Set. A, 266: 425-553. McCormick, M.P., Swissler, T.J., Chu, W.P. and Fuller, W.H., Jr., 1978. Post-volcanic stratospheric aerosol decay as measured by lidar. J. Atmos. Sci., 35: 1296--1303. Meinel, A.B., Meinel, M.P. and Shaw, G.E., 1976. Trajectory of the Mt. St. Augustine 1976 eruption ash cloud. Science, 193: 420--422. Moore, J.G., Nakamura, K. and Alcarez, A., 1966. The 1965 eruption of Taal volcano. Science, 151: 955--960. Morton, B.R., Taylor, G. and Turner, J.S., 1956. Turbulent gravitational convection from maintained and instantaneous sources. Proc. R. Soc., Set. A., 234: 1--23. Newhall, C.G. and Self, S., 1982. The volcanic explosivity index (V.E.I.): An estimate of explosive magnitude for historical volcanism. J. Geophys. Res., 87: 1231--1238. Pollack, J.B. and Ackerman, T.P., 1983. Possible effects of the'E1 Chichon volcanic cloud on the radiation budget of the northern tropics. Geophys. Res. Lett., 10: 1057--1060. Pollack, J.B., Toon, O.B., Sagan, C., Summers, A., Baldwin, B. and Van Camp, W., 1976. Volcanic explosions and climatic change: A theoretical assessment. J. Geophys. Res., 81: 1071--1083. Rose, W.I., Jr., 1972. Notes on the 1902 eruption of Santa Maria volcano, Guatemala. Bull. Volcanol., 36: 29--45. Settle, M., 1978. Volcanic eruption clouds and the thermal power output of explosive eruptions. J. Volcanol. Geotherm. Res., 3: 309--324. Simkin, T. and Howard, K.A., 1970. Caldera collapse in the Galapagos Islands, 1968. Science, 169: 429--437. Sparks, R.S.J. and Wilson, L., 1982. Explosive volcanic eruptions -- V. Observations of plume dynamics during the 1979 Soufri~re eruption, St. Vincent. Geophys. J.R. Astron. Soc, 69: 551--570. Thomas, G.E., Jakosky, B.M., West, R.A. and Sanders, R.W., 1983. Satellite limbscanning thermal infrared observations of the E1 Chichon stratospheric aerosol: First results. Geophys. Res. Lett., 10: 997--1000. Thorarinsson, S. and Sigvaldason, G.E., 1972. The Hekla eruption of 1970. Bull. Volcanol., 36: 269--288.
255 Toon, O.B. and Pollack, J.B., 1982. Stratospheric aerosols and climate. In: R.C. Whitten (Editor), The Stratospheric Aerosol Layer. Springer-Verlag, New York, pp. 121--147. Turco, R.P., 1982. Models of stratospheric aerosols and dust. In: R.C. Whitten (Editor), The Stratospheric Aerosol Layer. Springer-Verlag, New York, pp. 93--119. Varekamp, J.C., Luhr, J.F. and Prestegaard, K.L., 1984. The 1982 eruptions of E1Chichon volcano (Chiapas, Mexico): Character of the eruption, ash-fall deposits, and gas phase. J. Volcanol. Geotherm. Res., 23: 39--68. Walker, G.P.L., 1981. Generation and dispersal of fine ash and dust by volcanic eruptions. J. Volcanol. Geothermal Res., 11 : 81--92. Wilcox, R.E., 1959. Some effects of recent volcanic ash falls with especial reference to Alaska. U.S. Geol. Surv., Bull., 1028-N: 409--476. Wilson, L., Sparks, R.S.J., Huang, T.C. and Watkins, N.D., 1978. The control of volcanic column heights by eruption energetics and dynamics. J. Geophys. Res., 83: 1829-1836.