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the troposphere of about 37 kt/yr and 107 kt/yr, respectively, in 2008 (derived from
Figure 1
3, Montzka et al ., 2011 ). In comparison, Br emissions by quiescent
volcanic degassing and from smaller eruptive events (considering short- and
long-lived volcanogenic Br species) have been estimated as 3 to 40 kt/yr (Aiuppa
et al. , 2005 and references therein; see also Chapter 8 ). Ground-based and satellite
monitoring of recent eruptions at Soufrière Hills (Montserrat) 2002 and Kasatochi
(Aleutians) 2008 revealed 350 and 50
-
120 t/yr of Br injection into the troposphere
and lower stratosphere, respectively (Bobrowski et al ., 2003 ; Theys et al ., 2009 ).
In contrast, the recent study by Hörmann et al .( 2013 ) roughly quanti
-
ed a total
BrO mass up to 1 kt during the sub-Plinian eruption of Kasatochi (9
11 August
2008) using GOME-2 satellite data, which is in the range of Br masses from
CAVA eruptions ( Table 16.1 ). All CAVA eruptions considered in the present
study developed eruption columns of > 18 km altitude that reached well into the
stratosphere (Kutterolf et al ., 2008 ). In order to evaluate the potential effect on
ozone of such eruptions, we need to estimate the proportion of the originally
erupted halogens actually entering the stratosphere.
Different numerical models, considering microphysical and partly chemical
processes in an ascending eruption column, simulate that interaction with hydro-
meteors results in a partial scavenging of halogens, causing only a fraction of the
halogen mass emitted at the vent to actually reach the stratosphere. Tabazadeh
and Turco ( 1993 ) modelled that almost all halogen load is lost through scavenging
by condensed water before an eruption column reaches the stratosphere. More
complex plume model studies, in contrast, consider scavenging mainly by ice
formed in the ascending eruption columns and suggest that 10% to more than
25% of emitted halogens may reach the stratosphere (Textor et al ., 2003a , b ) . The
scavenging efficiency strongly depends on the composition of the gas phase,
including its salinity, the formation of ice crystals and on the dry versus wet nature
of the atmosphere (Textor et al ., 2003a ). Additionally, the properties of simultan-
eously emitted ash particles present in the atmosphere can in
-
uence the formation
of halogen-rich precipitates and aerosols from the gas phase and therefore change
the potential of eruption columns to deliver halogens into the stratosphere
(Delmelle et al ., 2005 ; Gislason et al ., 2011 ). Two contrasting scenarios for the
degree of preservation of halogens in the stratosphere have been observed: The
'
wet ' Pinatubo eruption 1991, which was comparatively halogen-rich (Shinohara,
2013 ), caused no measurable accumulation of stratospheric halogens since it
occurred simultaneously with the typhoon Yunya, probably leading to ef
cient
wash-out of volcanic halogens over a large region over the Philippine peninsula
Luzon on 15 June 1991 (McCormick et al ., 1995 ). On the other hand, the Hekla
2000 eruption plume, which reached the upper troposphere/lower stratosphere
at high latitudes (dry atmosphere, low tropopause height), generated local ozone
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