JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 10
PAGES 1721–1764
1998
Magmagenesis at Soufriere Volcano,
St Vincent, Lesser Antilles Arc
EMILY HEATH1, RAY MACDONALD1∗, HARVEY BELKIN2,
CHRIS HAWKESWORTH3 AND HARALDUR SIGURDSSON4
1
ENVIRONMENTAL SCIENCES DIVISION, IEBS, LANCASTER UNIVERSITY, LANCASTER LA1 4YQ, UK
2
US GEOLOGICAL SURVEY, RESTON, VA 22092, USA
3
DEPARTMENT OF EARTH SCIENCES, THE OPEN UNIVERSITY, MILTON KEYNES MK7 6AA, UK
4
GRADUATE SCHOOL OF OCEANOGRAPHY, UNIVERSITY OF RHODE ISLAND, NARRAGANSETT, RI 02882, USA
RECEIVED MAY 10, 1997; REVISED TYPESCRIPT ACCEPTED APRIL 16, 1998
Soufriere volcano of St Vincent (<0·6 Ma) is composed of basalts
and basaltic andesites, the most mafic of which (mg-number 75)
may be representative of the parental magmas of the calc-alkaline
suites of the Lesser Antilles arc. Parental, possibly primary, magmas
at Soufriere had MgO ~12·5 wt % and were probably nephelinenormative. They last equilibrated with mantle at ~17 kbar pressure,
at temperatures of around 1130°C and f(O2) exceeding FMQ
(fayalite–magnetite–quartz) +1. They fractionated, along several
liquid lines of descent, through to basaltic andesites and rarer
andesites over a range of crustal pressures (5–10 kbar) and
temperatures (1000–1100°C), separating initially olivine + Crspinel + clinopyroxene + plagioclase ± titanomagnetite and then
clinopyroxene + plagioclase + titanomagnetite + orthopyroxene
assemblages. The total amount of crystallization was some 76 wt
%. Amphibole was apparently not a fractionating phase. Sr and
Nd isotopic and trace element systematics show no evidence for
significant crustal assimilation. There is conflicting evidence as to
the pre-eruptive water contents of Soufriere magmas; compositions
of clinopyroxene phenocrysts and melt inclusions suggest H2O >3
wt %, whereas various projections onto phase diagrams are more
consistent with relatively anhydrous magmas. Primary magmas at
Soufriere were generated by around 15% melting of mid-ocean ridge
basalt type mantle sources which had been modified by addition of
fluids released from the slab containing contributions from subducted
sediments and mafic crust.
INTRODUCTION
KEY WORDS: high-MgO arc magmas; geochemistry; magmagenesis; Lesser
Antilles; Soufriere St Vincent
The Lesser Antilles intra-oceanic arc is a 750 km long
chain of volcanic islands resulting from the subduction
of rocks of Jurassic to Cretaceous age of the American
plate beneath the eastern edge of the Caribbean plate
(Fig. 1). Plate convergence rates are relatively slow; since
the Middle Eocene they have averaged 2·0–2·2 cm/year
(Pindell et al., 1988), compared with the arc average of
6·5 cm/year (Gill, 1981). Brown et al. (1977) showed that
the compositions of volcanic rocks vary along the active
arc, allowing the islands to be grouped according to three
magma series: tholeiitic in the islands north of Montserrat,
calc-alkaline in the central islands (Montserrat to St
Lucia), and alkaline in the southernmost islands (Grenada
and southern Grenadines). Those workers proposed that
the volcanic rocks of St Vincent are transitional, in terms
of magmatic affinity, between the southern and central
island suites, consistent with the geographical position of
the island. Thirlwall et al. (1994), on the other hand,
suggested that high-MgO basalts from St Vincent are
transitional between tholeiitic and calc-alkaline, and
Smith et al. (1996) referred to them as being tholeiitic.
Although recognizing the transitional nature of the suite,
we shall refer to the Soufriere rocks as calc-alkaline.
The occurrence of different magma series within a
close spatial and temporal context and the common
presence of magnesian (MgO >10 wt %) lavas in the
southern islands present a rare opportunity to assess the
factors which control magma compositions in a modern
intra-oceanic arc setting. The more magnesian members
∗Corresponding author. Telephone: 01524 593934. Fax: 01524 593985.
 Oxford University Press 1998
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 10
OCTOBER 1998
Fig. 1. Map of the Lesser Antilles island arc.
(MgO ~12·5 wt %) of the Soufriere St Vincent suite
represent (near-)primary magmas of a type which acted
as parental to the calc-alkaline suites of the central islands.
St Vincent is probably the only island in the arc where
such rocks are relatively common. We know of only
one other occurrence in the calc-alkaline suites, the
olivine–clinopyroxene-phyric picritic basalt of Ilet à
Ramiers, Martinique (Westercamp & Mervoyer, 1976).
There is a compositional continuum between magnesian basalts and andesites on Soufriere. This provides
an opportunity to contribute to the debate as to whether
generation of calc-alkaline series is by closed-system fractional crystallization of parental basalts or by combined
fractional crystallization and crustal contamination. Furthermore, we evaluate the role of water in the evolution
of the suite, particularly the issue of whether the mafic
magmas were water rich.
GEOLOGICAL HISTORY
Figure 2 is a generalized geological map of St Vincent;
no detailed map exists. The Soufriere stratovolcano dominates the northern half of the island and is the most
active subaerial volcano in the arc. K–Ar dating has
placed a lower limit of 0·6 Ma on the age of Soufriere
(Briden et al., 1979). There are several other major
volcanic centres on the island which are no longer active;
the ages of the Richmond Peak–Mt Brisbane centres
to the immediate south of Soufriere, and the Grand
Bonhomme centre further south are not precisely known,
but the pre-Soufriere lavas dated by Briden et al. (1979)
yielded K–Ar ages of between 1 and 3 Ma.
The geological evolution of the volcano has been
characterized by four main volcanic formations (Sigurdsson & Carey, 1991). These represent protracted
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MAGMAGENESIS AT SOUFRIERE VOLCANO
Fig. 2. Generalized geological map of northern St Vincent (after Rowley, 1978).
periods of predominantly effusive or explosive volcanism,
which may be controlled more by the geomorphological
and hydrological features of the volcano than by the
silica and volatile contents of the magmas (Sigurdsson
& Carey, 1991). Early activity (~0·6 Ma–10 ka) was
characterized by the extrusion of basaltic and basaltic
andesite lavas from a central vent, with <5% pyroclastic
deposits. These are named the Pre-Somma Lavas because
they pre-date a probable major structural failure of the
volcano’s southern flank which created the Somma scarp
and generated a thick debris flow. Some Pre-Somma
lavas from the southern flanks of Soufriere are unusually
magnesian, and may possibly have been erupted from
a different centre or centres. A series of well-bedded
pyroclastic fall deposits mantles much of the island of St
Vincent, and has been tentatively correlated with coarse
tephra beds exposed in the crater and on the flanks of
Soufriere. The units range from black scoria to yellow
lapilli and pumiceous tuff, and are collectively known as
the Yellow Tuff Formation, with an estimated volume
of 48 km3 (Rowley, 1978). Carbon dating indicates that
the formation spans the period 3600–4500 years bp. The
rarity of unconformities suggests rapid deposition. The
vents feeding the Yellow Tuff Formation have not been
identified with certainty; but the sequence was probably
erupted from the central Soufriere vent (Sigurdsson &
Carey, 1991).
A predominantly effusive phase of activity followed the
emplacement of the Yellow Tuff Formation, with the
eruption of ponded basaltic and basaltic andesitic Crater
Lavas. The most recent phase of activity has been characterized by vulcanian explosive eruptions, generating a
thick succession of pyroclastic fall and flow deposits,
named the Pyroclastic Formation. New 14C dates (Sigurdsson et al., 1998) suggest that the Pyroclastic Formation may extend back much further than the
1300 years indicated by Sigurdsson & Carey (1991), and
possibly overlaps with the Yellow Tuff Formation. There
have been at least five major historic eruptions of the
Soufriere (1718, 1812, 1902, 1971, 1979); the activity
has been characterized by the extrusion of basaltic andesite lava domes in the crater area followed by phreatomagmatic explosions generating pyroclastic flows.
PETROGRAPHY, MINERALOGY AND
PHYSICAL PARAMETERS
Petrography
Petrographic descriptions of Soufriere eruptive rocks have
been given by Roobol & Smith (1975), Carey & Sigurdsson (1978), Shepherd et al. (1979), Graham &
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Thirlwall (1981), Devine & Sigurdsson (1983), Dostal et
al. (1983) and Bardintzeff (1984, 1992). Most of these
studies concentrated on the products of the activity since
1902.
Detailed sample and locality descriptions of rocks used
in this study are given in Appendix A. Soufriere is
formed almost entirely of basalts and basaltic andesites
[classification scheme of LeBas et al. (1986)]; andesites
are found only as components of mixed magma rocks of
the Yellow Tuff Formation. Both basalts and basaltic
andesites have been erupted throughout the volcano’s
history, although basalts were volumetrically at their most
abundant in the earlier stages (Pre-Somma and Yellow
Tuff Formations). Roobol & Smith (1975) recorded a
progressive change from early erupted basaltic andesite
to late basalt in the 1902–1903 activity, perhaps pointing
to the existence of zoned magma chambers at least
spasmodically beneath Soufriere.
Table A1 shows the modal proportions (volume percent, recalculated vesicle- and groundmass-free) of phenocrysts in representative samples from each of the main
geological formations of Soufriere. The basalts range
from microphyric, fine-grained rocks with abundant (up
to 30%) microphenocrysts of ol + spinel ± cpx, to
more coarsely porphyritic rocks also containing phyric
plagioclase. Basaltic andesites are generally phenocryst
rich (35–60%), containing the assemblage cpx + Timag + plag + opx ± ol. The order of appearance of
phases was ol + Cr-sp ± Ti-mag, followed by cpx, plag
and then opx.
Although it forms a core to an augite crystal in STV
303, pigeonite occurs mainly as thin rims around orthopyroxene, clinopyroxene and olivine crystals in the
Pre-Somma Lavas (STV 315, 318, 323). Titanomagnetite
microphenocrysts are present in some basalts but are
always subordinate to Cr-spinel. Ilmenite is present only
as a groundmass phase. Amphibole has not been recorded
as a phenocryst at Soufriere, though Shepherd et al.
(1979) and Graham & Thirlwall (1981) found corroded
amphiboles in products of the 1979 eruption which were
assumed to be xenocrystic. We have found rounded
crystals of amphibole in STV 354 (from the 1979 eruption) and STV 363 (scoria of unknown age from the
Pyroclastic Formation). The amphiboles are 2 mm long,
greenish yellow in colour, and contact plagioclase crystals.
The aggregates probably represent parts of cumulate
blocks. Only one occurrence of phyric apatite has been
found, as a microphenocryst included in a plagioclase
phenocryst in andesite STV 376(L).
There is a marked tendency for the phenocryst phases
to form clusters, up to 2 mm across. These vary from
monomineralic, clinopyroxenitic clots through olivine–
clinopyroxene-rich clusters to titanomagnetite-rich gabbros. STV 358 contains a 4 mm × 3 mm olivine
NUMBER 10
OCTOBER 1998
aggregate where the crystals are strained and partially
recrystallized.
Melt (glass) inclusions occur in all the phenocryst phases
(Graham & Thirlwall, 1981; Devine & Sigurdsson, 1983;
Bardintzeff, 1992). Those in olivine tend to be large
(50–100 lm) and spheroidal, and generally have large
contraction bubbles. There is a tendency for inclusions
in olivines in the most magnesian basalts to be at least
partly devitrified. Inclusions in orthopyroxene are more
abundant and larger (Ζ50 lm) than those in clinopyroxene (Ζ30 lm), whereas in plagioclase large (Ζ100
lm) subrectangular to spheroidal inclusions form trains
in crystal cores.
Macroscopic evidence (e.g. banded and mingled pumices) for mixing between basaltic andesite and dacite at
Soufriere has been documented for the 1902 (Carey &
Sigurdsson, 1978) and 1979 (Shepherd et al., 1979; Graham & Thirlwall, 1981; Devine & Sigurdsson, 1983;
Bardintzeff, 1992) eruptions. Light-coloured bands of
dacite in scoria blocks in the 1979 ejecta contain feldspar,
orthopyroxene, quartz and magnetite crystals, and some
contain partially fused granitic xenoliths (Graham &
Thirlwall, 1981; Devine & Sigurdsson, 1983). They are
evidence of at least local crustal fusion beneath Soufriere.
We have also found clear evidence of mixing between
basalt and andesite and between basaltic andesite and
andesite in western outcrops of the Yellow Tuff Formation
and in scoria from the Pyroclastic Formation.
Cumulate-textured blocks are found at most volcanoes
in the Lesser Antilles, but are particularly abundant and
well documented at Soufriere, St Vincent. Mineralogy
and textures are highly variable, but hastingsitic amphibole, calcic plagioclase, olivine, titanomagnetite and
high-Ca pyroxene are common cumulate phases (Lewis,
1973a, 1973b; Arculus & Wills, 1980; Dostal et al., 1983).
Metavolcanic and calc-silicate sedimentary xenoliths
are also common at Soufriere (Devine & Sigurdsson,
1980; Carron & Le Guen de Kerneizon, 1991). We have
analysed five metamorphic xenoliths to help constrain
the nature of any potential crustal assimilant; the protoliths were tuffs (STV 336, 340), basalt (STV 353), and
andesite (STV 337, 339). STV 337 is cordierite bearing.
Phenocryst compositions
In this section, we refer to various mineral–melt partitioning data, using whole-rock compositions as proxies
for melt compositions. We appreciate that this creates
problems in strongly porphyritic rocks, in that later stages
of phenocryst growth may have been from melts widely
removed in composition from the bulk rock. We have
tried to take account of this effect by using only phenocryst
core compositions, which we assume to have crystallized
close to the liquidus. Mineral compositional data for
Soufriere samples are available from R.M. on request.
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MAGMAGENESIS AT SOUFRIERE VOLCANO
Olivine phenocrysts in the most magnesian basalts (e.g.
STV 301) take three forms: (1) euhedral to subhedral
prisms up to 1 mm in size, with core compositions ranging
from Fo89 to Fo87 and substantial zoning (up to 19% Fo);
(2) smaller (<0·5 mm) rounded crystals, with slightly more
Fe-rich cores (Fo86–83) and zoning <14% Fo; (3) rare,
ragged, variably resorbed, probably xenocrystic, grains,
with cores of Fo80 and only weak, normal zoning (~3%
Fo). No reverse zoning has been found; this suggests that
crystallization occurred in a system where there was
no, or minimal, input of fresh magma. In the basaltic
andesites, olivines are small (<1 mm) and commonly
embayed through resorption; mantling textures indicate
an olivine–orthopyroxene reaction relationship. Core
compositions range from Fo85 to Fo55, with rare values
of Fo90. NiO contents reach 0·38% and are positively
correlated with mg-number.
The most magnesian olivines (mg-number ~89) are
found as phenocrysts in the most magnesian basalts, e.g.
STV 301. This rock has an mg-number of 73 [using an
Fe2O3/FeO weight ratio of 0·22, derived from the Kilinc
et al. (1983) expression for Fe speciation, and an f (O2)
of FMQ (fayalite–magnetite–quartz) + 1·7 (see below)].
Olivine of composition Fo89 could have crystallized from
melt with mg-number 73 if KD = 0·33. This value is the
same as that determined by Wagner et al. (1995) for
olivines grown in 1 kbar, water-saturated experiments
on a high-Al2O3 basalt from the Medicine Lake volcano,
California, and similar to the average value of 0·34
determined for the 1 kbar water-saturated experiments of
Sisson & Grove (1993b). Ulmer (1989) has demonstrated
experimentally that Mg–Fe2+ partitioning between olivine and liquid in a calc-alkaline picrobasalt is pressure
dependent, and reported a range of KD from 0·315 at 1
bar to 0·365 at 25 kbar. The value for 15 kbar, close to
the possible equilibration pressure of 17 kbar for STV
301 (see Figs 8 and 9, below) was 0·341. We conclude,
therefore, that STV 301 represents the most primitive
melt erupted by Soufriere. Even more primitive rocks
have been recorded from other arcs; Eggins (1993) has
collated compositions of olivine phenocrysts from several
arc systems, some as high as Fo94.
On plots of Fe2+/Mg ratios of olivine phenocrysts
against whole rocks (Fig. 3a), core compositions, in particular, plot on both sides of any reference KD line, that
is, they apparently have compositions that are either too
evolved (above the line), or too primitive (below) for
the whole-rock composition. There are several possible
explanations for the scatter:
(1) The P–T–X dependence of KD (Ulmer, 1989) means
that KD should constantly change to lower values as
magma evolves, and not remain constant as is commonly
assumed.
(2) Relatively primitive core compositions in basaltic
andesites, e.g. Fo86 in SVE 113 (SiO2 54·8 wt %) may
represent phenocrysts which did not re-equilibrate from
higher-temperature stages of magma evolution, and/or
the products of mixing with a more primitive basaltic
magma. We note, however, that reverse zoning is uncommon in Soufriere phenocrysts.
(3) The compositions of the rims of many olivines are
more Fe rich than could be predicted from equilibrium
relationships (Fig. 3a). This again may reflect magma
mixing, this time with a more evolved melt, or differential
amounts of re-equilibration of olivine microphenocrysts
by solid-state diffusion at magmatic temperatures.
In all rock types, a spinel phase forms small (p0·2 mm)
euhedral to subhedral inclusions in the cores of olivine
and, less frequently, pyroxene phenocrysts, and also
occurs more rarely as partially resorbed, discrete microphenocrysts, the size increasing from <0·5 mm in basalts
to 0·8 mm in basaltic andesites. There is a continuum of
compositions from Cr-spinel to titanomagnetite, although
the majority of analyses tend to be bimodal (Fig. 4); in
Cr-spinel, cr-number ranges from 35 to 85 (the majority
>50), mg-number from 12 to 54 and Fe3+/
(Fe3+ + Al + Cr) from 10 to 80. TiO2 varies from 0 to
10 wt %. In titanomagnetites, the same ranges are 0–35
(most <10), 5–26 and 81–98, respectively, and TiO2
varies from 5 to 30 wt % (Fig. 4). Compositional zoning
has been measured in only a handful of larger microphenocrysts and shows Mg/Fe, Cr/Al and Fe2+/Fe3+
ratios decreasing, and TiO2 abundances increasing, towards crystal rims. Differentiation within the high-temperature spinels led to ferrite and ulvöspinel enrichment
with little change in cr-number (Fig. 4a). This is typical
of crystallization under relatively oxidizing conditions
(Ballhaus et al., 1991). There is a crude, positive correlation between Fe2+/Mg ratios in spinels and whole
rocks, suggesting an approach to equilibrium crystallization.
The Soufriere spinels are similar to spinel compositions
recorded from island arcs elsewhere; on a Cr–Al–Fe3+
plot (Fig. 4b), for example, they largely fall within the
fields of arc basalts constructed by Eggins (1993). The
compositional range, and particularly the continuum
between Cr-spinels and titanomagnetites, closely matches
spinels from picrites of Ambae volcano in the Vanuatu
arc (Eggins, 1993).
The majority of clinopyroxene phenocrysts are augite,
although diopsidic cores are found in olivines and augites
in more magnesian rocks, e.g. STV 301 and STV 334.
They are typically 0·5–2 mm in size, subhedral, with
weak zoning, partially resorbed margins and inclusions
of Fe–Ti–Cr oxides. There is a strong tendency in some
rocks for pyroxene and olivine crystals to form clusters
2 mm in diameter.
Maximum Cr concentrations (Ζ0·9 wt %) are found
in clinopyroxene cores in some more magnesian basalts
(STV 301, 334); values in more evolved rocks are typically
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Fig. 3. (a–c) Average Fe2+/Mg in cores and rims of olivine, clinopyroxene and orthopyroxene phenocrysts, respectively, plotted against Fe2+/
Mg in host rocks. Continuous lines are for equilibrium KD values. (d) Plot of average Ca/Na (cations) in cores and rims of plagioclase phenocrysts
and in host rocks. Continuous lines represent approximate exchange KD, which show a progressive increase with melt H2O content from 1·0 for
anhydrous melts through ~1·7 for melts with 2 wt % H2O and 3–4 for melts with 4 wt % H2O to 5·5 for melts with 6 wt % H2O (Sisson &
Grove, 1993a, fig. 1).
<0·1 wt %. The Soufriere clinopyroxenes are notably Al
rich, with maximum a.f.u. values >0·25 (Al2O3 > 6 wt
%) in some basalts. Variation in Al with mg-number in
the clinopyroxenes (Fig. 5a) shows a diffuse peak at mgnumber 80. This is similar to the situation in ankaramites
from western Epi (Barsdell & Berry, 1990) and in picrites
from Ambae (Eggins, 1993), both in the Vanuatu arc.
In all three cases, the peak coincides with the commencement of plagioclase crystallization. The clinopyroxenes in Soufriere basalts reach relatively high AlVI
values (Ζ0·13 a.f.u.). Although such high values might
indicate high-pressure crystallization (Kennedy et al.,
1990), the fact that the distributions of AlIV and AlVI as
a function of mg-number are closely similar to that of
RAl suggests that the AlIV/AlVI ratio is composition
dependent. Ti behaviour in Soufriere clinopyroxenes
(Fig. 5b) fairly closely mimics that of Al, as it does in the
Ambae suite (Eggins, 1993), though the abundances in
Soufriere basalts tend to be higher.
Compositional zoning is most commonly normal, with
rimwards enrichment in Fe (<5% Fs) and decrease in
Cr, Al and sometimes Ti. Reverse zoning is restricted to
some basaltic andesites and is typically at the limit of
analytical resolution (Fs Ζ1%).
Clinopyroxene and whole-rock compositions correlate
fairly well, suggesting that quasi-equilibrium conditions
prevailed during clinopyroxene crystallization (Fig. 3b).
Relationships are consistent with an equilibrium KD
around 0·4 in the basalts and 0·35 in more evolved rocks.
These values are higher than those (~0·20–0·25) found
experimentally in 1 atm, anhydrous experiments on midocean ridge basalt (MORB) compositions and silicasaturated and -undersaturated arc lavas (Grove & Bryan,
1983; Grove & Baker, 1984; Kennedy et al., 1990) and
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HEATH et al.
MAGMAGENESIS AT SOUFRIERE VOLCANO
Fig. 4. (a) Variation of cr-number [100 Cr/(Cr + Al)] with mg-number [100 Mg/(Mg + Fe2+)] in Cr-spinels and titanomagnetites present as
phenocrysts and inclusions in Soufriere rocks; (b) Cr–Al–Fe3+ variation in Cr-spinel and titanomagnetite phenocrysts and inclusions in Soufriere
rocks, compared with fields for spinels in island arc basalts (IAB) presented by Eggins (1993, fig. 6).
1 kbar water-saturated experiments by Sisson & Grove
(1993b), but comparable with that of 0·38 determined in
an Aleutian high-magnesia basalt at 12 kbar, 1315°C
under anhydrous conditions by Johnston & Draper
(1992). A possible explanation for the elevated KD values
is the relatively aluminous nature of the Soufriere clinopyroxenes. Sisson & Grove (1993a) found experimentally
that Al-rich sectors of zoned clinopyroxenes synthesized
from natural high-alumina basalts had higher KD (up to
0·31) than less aluminous sectors.
Orthopyroxene crystallized instead of pigeonite in the
basaltic andesites presumably because either (1) the melt
temperatures were too low for the likely equilibration
pressures ([5 kbar) and the mg-number of the rocks
(Fig. 6), and/or (2) the f (O2) of the magmas was too
high. Pigeonite (Wo7·8 En65·5 Fs26·7) has been found as a
core in an augite crystal in STV 303 which may have
been relatively water poor and more reduced. We ascribe
the occurrence of pigeonite rims (Wo5–11 En48–66 Fs28–45)
to the other mafic phases to crystallization during
magma ascent, as the pigeonite–orthopyroxene inversion
is shifted to lower temperatures at lower pressures
(Lindsley, 1980).
Plagioclase phenocrysts are abundant in all but the
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Fig. 5. Variations in (a) Al and (b) Ti cations per six oxygens as a function of mg-number in clinopyroxene phenocrysts.
Fig. 6. Temperature plotted against mg-number showing the experimentally determined stability fields of pigeonite and orthopyroxene ( + augite)
at pressures of 1·5 GPa and 0·1 MPa (after Kersting & Arculus, 1994, fig. 11). Temperatures for Soufriere rocks taken from Table 2; diamonds
are basalts, triangles are basaltic andesites. Data for the Klyuchevskoy suite, Kamchatka, are shown for comparison (dotted box; Kersting &
Arculus, 1994).
most magnesian rocks, ranging in size from 5 mm to
microphenocrysts <0·5 mm. Phenocrysts are commonly
euhedral, with conspicuous normal or fine (~1 lm)
oscillatory zoning. Compositions range from An97 to An46;
anorthitic cores are common (An >90), whereas rims
tend to be more sodic (An70–50). Average core compositions
are shown in Fig. 3d but the range of core compositions
in individual specimens can exceed 30% An. The extent
of zoning in single grains may exceed 40% An. No
reverse zoning has been recorded. Plagioclase inclusions
are frequently found within olivine and pyroxene phenocrysts, and these exhibit an even greater range of
compositions (An96–An8) than discrete phenocrysts. The
most calcic compositions are from inclusions in olivine,
whereas the most sodic plagioclases usually occur included in clinopyroxenes.
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MAGMAGENESIS AT SOUFRIERE VOLCANO
Ca/Na ratios in plagioclase cores and rims are plotted
against whole-rock Ca/Na ratios in Fig. 3a. The spread of
core compositions at given whole-rock Ca/Na probably
reflects the presence of unre-equilibrated crystals from
higher-temperature stages of evolution, mixing between
more and less evolved magmas, and crystallization of
plagioclase from variably water-rich magmas. Sisson &
Grove (1993a), for example, have shown experimentally
that KD at 2 kbar varies with water content of the melt,
from 1·0 for anhydrous melts to 5·5 for melts with 6%
water (Fig. 3d).
Orthopyroxene occurs in Soufriere rocks with >51%
SiO2 as (1) small (<2 mm), pink, euhedral to subhedral
prisms, (2) anhedral crystals in pyroxene–olivine clusters,
(3) reaction rims around olivine, or more rarely clinopyroxene, crystals, and (4) cores to clinopyroxene crystals.
The total compositional range is En70–54. Crystals are
only weakly, normally zoned, rimward decreases in En
usually being Ζ2%. TiO2 and Al2O3 tend to increase
rimwards.
The relationship between orthopyroxene and wholerock Fe2+/Mg ratios (Fig. 3c) is consistent with crystallization of orthopyroxene under close to equilibrium
conditions and with KD somewhat greater than 0·4. This
is comparable with the value of 0·37 determined for
orthopyroxene in a high-magnesia basalt from Umnak
island in the Aleutian arc at 12 kbar, 1315°C under
anhydrous conditions by Johnston & Draper (1992). Such
values are higher than the 0·27 used, for example, by
Hunter & Blake (1995), and may, as suggested above for
the clinopyroxenes, be related to the relatively aluminous
nature of the Soufriere orthopyroxenes (Al2O3 Ζ2·6 wt
%).
Intensive parameters
Barometry
There is currently no reliable mineral geobarometer
which allows us to estimate total lithostatic pressure
without an independent assessment of magmatic water
contents. We shall return to the question of water contents
later. Here we estimate lithostatic pressure using two,
experimentally determined, sets of phase relationships,
at least one of which is apparently only weakly dependent
on the water content of the system.
Sobolev & Danyushevsky (1994) used the Ol–Q–Pl
projection (from Di) of Walker et al. (1979) and calibrated
ol + opx ± cpx cotectics to determine the pressure of
the last equilibration of primary melts with their mantle
sources (Fig. 7). The most primitive Soufriere rocks
[(ol + sp ± cpx)-phyric] lie close to an interpolated 17
kbar cotectic for both dry and wet (2 wt % H2O) melts,
which, if we assume that these rocks were close to
saturation with orthopyroxene at higher pressures, corresponds to an equilibration depth of ~55 km. Given a
crustal thickness of ~30 km beneath St Vincent (Boynton
et al., 1979), this implies equilibration within the uppermost 30 km or so of the mantle. Rocks in which
plagioclase had apparently just become a liquidus phase
plot in the range 17–14 kbar. More evolved rocks,
including the most mafic opx-phyric specimen (STV 315)
plot down-pressure from the 10 kbar cotectic, in an
uncalibrated region of the diagram but implying equilibration within the crust.
The pressure at which the (olivine ± clinopyroxene)phyric basalts equilibrated with mantle peridotite can
also be estimated using the system diopside
(Di)–jadeite + calcium
Tschermak’s
molecule
( Jd + CaTs)–olivine (Ol)–quartz (Qz). Figure 8 is a
projection from Di into the plane ( Jd + CaTs)–Ol–Qz,
on which are shown cotectics, defined by experiments
on the Tinaquillo lherzolite, representing the locus of
liquids in equilibrium with ol + opx + cpx + sp over
the pressure range 0·5–4·0 GPa (after Falloon et al.,
1988). The most magnesian rocks project to a pressure
of ~1·7 GPa (17 kbar). More evolved basalts plot at
lower pressures; orthopyroxene apparently became a
liquidus phase at ~0·5 GPa (5 kbar). Despite the fact
that the cotectics shift away from olivine in the hydrous
system, the pressure estimates are similar to those from
the previous projection.
The ~17 kbar pressure estimates for the (olivine ±
clinopyroxene)-phyric basalts refer to the final pressure
of equilibration with mantle peridotite. It is possible
that even the most primitive Soufriere magmas were
themselves derived from more olivine-rich parents; for
example, Thirlwall et al. (1996) have inferred that the
parental magmas to the basaltic suites of Grenada were
picritic basalts with MgO contents of ~16 wt %. Such
parental magmas may have been an integration of melt
fractions or batches from various levels in the mantle
wedge. There may even have been mantle fusion at
these relatively shallow depths, especially if the isotherms
beneath St Vincent have been perturbed upwards by the
passage of magma batches through reused conduits, in
the manner proposed for Umnak island in the Aleutians
by Johnston & Draper (1992).
Thermometry
Previous temperature estimates for Soufriere rocks are
summarized in Table 1. There are some inconsistencies
in these estimates, particularly the fact that the basaltic
andesites yield higher temperatures than the basalts.
We have used, therefore, the olivine–spinel exchange
thermometer of Ballhaus et al. (1991) and the two-pyroxene thermometer of Lindsley (1983) to try to determine
more precisely the crystallization temperatures of Soufriere rocks. Unfortunately, the absence of phenocrysts of
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Fig. 7. Projection from Di into the Ol–Pl–Qz plane of the system Ol–Di–Pl–Qz, using the algorithm of Walker et al. (1979). Dashed lines are
cotectics, for pressures in the range 10–30 kbar, for melts in equilibrium with ol + opx ± cpx under conditions of anhydrous melting of the
mantle; continuous lines are the approximate positions of these cotectics for 2 wt % H2O in the melt (after Sobolev & Danyushevsky, 1994, fig.
9). Soufriere rocks are distinguished by phenocryst assemblage.
ilmenite precluded use of the coexisting oxides geothermometer.
Successful application of the Ballhaus et al. (1991)
thermometer relies, inter alia, on the correct choice of (1)
equilibration pressure, (2) olivine–spinel pairs coexisting
in equilibrium, and (3) oxidation ratio of the spinel phase.
We have used pressures estimated from Figs 7 and 8
(ranging from 1·7 GPa for STV 301 to 0·5 GPa for
STV 315), and oxidation ratios based on stoichiometry.
Furthermore, we have used only basalts (i.e. SiO2 <52
wt %), as olivine is in reaction relationship with orthopyroxene in the basaltic andesites. The most magnesian
olivine and Cr-spinel compositions were used, but the
results were found to be extremely sensitive to slight
variations in input compositions (particularly of olivine).
To illustrate this, a range of temperatures (and oxygen
fugacities) are shown in Table 2, corresponding to a
range of olivine mg-numbers. It is usually difficult to
determine simply on textural grounds which highly magnesian olivine composition most closely approaches equilibrium with respect to the most magnesian Cr-spinel
inclusion found in each rock.
Calculated temperatures range from 1026 to 1147°C
(Table 2). The higher end of this range is more likely to
represent realistic crystallization temperatures for basalts,
and compares well with previously determined temperatures for Soufriere rocks (Table 1).
We applied the two-pyroxene thermometer to the nine
samples for which we have data for coexisting pyroxenes.
In most samples, the pyroxenes apparently coexisted
close to equilibrium, in the sense that tie-lines do not
cross (Fig. 9) and mg-numbers are similar. Only four,
however, gave consistent temperatures for both pyroxenes, in the sense that the ranges for each pyroxene
overlap. All four are basaltic andesites within the narrow
compositional range SiO2 54–56%. Temperatures (for
pressures of 5 kbar) are ~1050–1060°C (±50°C), a
little lower than most of those quoted in Table 1, and
overlapping with the temperature ranges for the basalts
which were estimated using the olivine–spinel geothermometer.
Oxygen fugacity
f (O2) has been estimated using the oxygen geobarometer
of Ballhaus et al. (1991), and the results are presented in
Table 2. Temperatures were taken from Table 2 and
pressures were inferred from Figs 8 and 9. We use only
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MAGMAGENESIS AT SOUFRIERE VOLCANO
Fig. 8. Soufriere rocks (crosses) projected from Di into the plane Ol–Qz–( Jd + CaTs). The continuous line represents the locus of liquids in
equilibrium with olivine only, and dashed lines show ol + opx ± cpx cotectics for anhydrous melting of the Tinaquillo lherzolite (TL; representing
mantle peridotite) at various pressures between 0·5 and 4·0 GPa (5–40 kbar), after Falloon et al. (1988).
Table 1: Previously published temperature estimates for Soufriere rocks
Rock type
Formation
T (°C)
Method
Ref.
B, BA
Unspecified
1130 ± 30
2-pyroxene (Wood & Banno, 1973)
1
B, BA
Unspecified
1000–1140
Plagioclase–melt (Kudo & Weill, 1970)
1
BA
1979 eruption
1100–1060
Fusion of devitrified melt inclusions
2
BA
1979 eruption
1165 ± 18
2-pyroxene (Wood & Banno, 1973)
3
BA
1979 eruption
1180–1120
Plagioclase–melt (Kudo & Weill, 1970)
3
BA
1979 eruption
~1170
Two-liquid consolution curve
4
Rock type: B, basalt; BA, basaltic andesite. References: 1, Dostal et al. (1983); 2, Devine & Sigurdsson (1983); 3, Bardintzeff
(1984); 4, Martin-Lauzer et al. (1986).
basalts, as the barometer is most applicable to mafic
rocks. No correction was made for the absence of orthopyroxene in the basalts, but the correction is small
for rocks of this composition (a few tenths of a log unit;
Ballhaus et al., 1991).
Using the Ballhaus et al. (1991) formulation, f (O2)
values in the basalts range from FMQ + 0·9 to + 1·9.
Such values are comparable with those estimated from
coexisting oxides in rocks from Statia and from Mt
Pelée, Martinique (Smith & Roobol, 1990). They are
also comparable with f (O2) estimates for olivine–spinel
pairs in other island arc lavas (e.g. FMQ + 1 to + 3;
Eggins, 1993), which are generally more oxidizing than
values obtained for MORB and intra-plate basalts (FMQ
–1 to FMQ + 1; Ballhaus et al., 1991). This suggests that
the mantle wedge over subduction zones is more oxidizing
than the mantle sources of other basalt types, perhaps
as a result of introducing fluids derived from the subducting slab into the wedge (Ballhaus et al., 1991; Wood,
1991).
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Table 2: T-f(O2) estimates for Soufriere rocks
Sample no.
Rock type
T (°C) (Lind.)
T (°C) (Ball.)
f (O2) (Ball.) FMQ+
mg-no. of ol used in
Ball. formulation
STV 301
B
—
1026–1130
1·8–1·5
88·5–86·5
STV 334
B
—
1089–1132
1·9–1·8
88·5–87·7
STV 315
B
—
1027–1147
1·3–0·9
88·5–86·2
STV 79-70
BA
1063
—
—
—
SVE 113
BA
1048
—
—
—
STV 318
BA
1043
—
—
—
STV 323
BA
1038
—
—
—
Rock type: B, basalt; BA, basaltic andesite. (Ball.) refers to temperatures and f (O2) calculated using the olivine–spinel
formulations of Ballhaus et al. (1991). (Lind.) refers to temperatures calculated using the two-pyroxene thermometer of
Lindsley (1983).
Fig. 9. Tie-lines (dashed) connect compositions of coexisting clinopyroxene and orthopyroxene phenocrysts in Soufriere rocks. Temperature
contours are from Lindsley (1983) and are for 5 kbar, considered from the pressure range in that paper to be the most appropriate pressure for
Soufriere basaltic andesites.
GEOCHEMISTRY
Analytical data for Soufriere rocks have previously been
presented by Tomblin (1968), Baker (1972), Pushkar et
al. (1973), Roobol & Smith (1975), Brown et al. (1977),
Rowley (1978), Graham & Thirlwall (1981), Devine &
Sigurdsson (1983), Dostal et al. (1983), White & Patchett
(1984), White & Dupré (1986), Bardintzeff (1992),
Thirlwall et al. (1994) and Turner et al. (1996). Our
data set essentially covers the compositional range as
established in these papers and for the sake of internal
consistency we use, with some exceptions (specified in
the text), only our data in subsequent sections. We present
major and X-ray fluorescence (XRF) trace element data
for 49 eruptive rocks and five metamorphic xenoliths
(Tables A2 and A3), instrumental neutron activation
analysis (INAA) data for 18 eruptive rocks (Table A4),
Sr isotopic ratios in 19 eruptive rocks and five xenoliths,
Nd isotopic ratios in 13 eruptive rocks and two xenoliths,
and Pb isotopic ratios in 14 eruptive rocks (Table A5).
Analytical techniques are described in Appendix B.
Liquid compositions?
Before the geochemical evolution of Soufriere can be
discussed, it is necessary to assess the extent to which
whole-rock compositions represent liquid compositions.
Modal proportions of individual phenocryst phases are
as high as 47 vol. % (e.g. plagioclase in STV 315; Table
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MAGMAGENESIS AT SOUFRIERE VOLCANO
A1), which might reflect accumulation. If such porphyritic
rocks represent liquid compositions, they must preserve
the equilibrium proportions of crystallizing phenocryst
phases. Furthermore, there is abundant evidence for
disequilibrium conditions during crystallization (complexly zoned phenocrysts, resorbed mineral textures).
Despite the large range in mineral compositions observable in individual rocks and between different rocks,
the phenocryst assemblages and average compositions
vary with whole-rock composition in a fairly predictable
manner and tie-lines between the coexisting mineral
phases generally do not cross.
Whereas Crawford et al. (1987) have suggested that
arc basalts with >18 wt % Al2O3 form by plagioclase
accumulation, Sisson & Grove (1993a) have shown that
liquids saturated with ol + pl + cpx + H2O at 2 kbar
contain ~20 wt % Al2O3. A 20 wt % value has been
used by Davidson et al. (1993) for identifying plagioclase
accumulation in Lesser Antilles rocks. At Soufriere, Al2O3
concentrations increase overall with increasing differentiation, as marked by decreasing MgO abundances
(see Fig. 14a, below), but not beyond the 20 wt %
threshold value. The lack of a pronounced positive Eu
anomaly in chondrite-normalized REE diagrams (see
Fig. 15, below) also argues against significant accumulation of plagioclase.
As for the possibility of olivine accumulation in some
of the most magnesian lavas, we noted earlier that the
olivine in the most primitive basalt (STV 301) was
apparently in, or close to, equilibrium with the wholerock composition. Furthermore, there are no simple
correlations in the basalts between modal olivine or
clinopyroxene and mg-number or Ni and Cr contents in
the whole rocks. There seems to be no good evidence
for olivine accumulation, a conclusion also reached for
C-series lavas of Grenada by Thirlwall & Graham (1984).
Finally, the compositions of melt inclusions in phenocrysts fairly closely match those of the whole rocks, at
least in the range covered by the whole rocks; Fig. 10
shows CaO–SiO2 relationships as an example. We suggest, therefore, that the Soufriere bulk rocks represent
(near-)liquid compositions.
silica-undersaturated mafic end-members of the M- and
C- series of Grenada (Fig. 11). The slope of the Soufriere
array in Fig. 11 is related to fractionation of silicaundersaturated cpx + spinel, in addition to ol + plag.
When plotted in standard discrimination diagrams
(Fig. 12), the St Vincent suite is not unequivocally assigned
to either tholeiitic or calc-alkaline category. Miyashiro
(1974) used FeO∗/MgO–SiO2 relationships to discriminate between the two series. He proposed that each
series should have progressed into the relevant field at
intermediate degrees of differentiation (i.e. 2 < FeO∗/
MgO < 5), that is, the discriminating feature is the slope
of the compositional trend. The Soufriere suite would be
more likely to be classified as tholeiitic on this basis,
though the data spread across the boundary (Fig. 12a).
On the K2O–SiO2 plot (Fig. 12b; after Rickwood, 1989)
the rocks scatter across the boundaries between the two
magma series. In an AFM diagram (Fig. 12c; after Irvine
& Baragar, 1971), the mafic members of the suite plot
in the tholeiite field, whereas more evolved members
trend into the calc-alkaline.
We feel that trace element and isotope data provide a
rationale for selecting calc-alkaline affinity for the Soufriere rocks. Figure 13, for example, presents MgO–(87Sr/
86
Sr)i relationships for low-K tholeiites, calc-alkaline and
C- and M-series rocks from the arc. The Soufriere data,
with almost constant 87Sr/86Sr and a large range in MgO
contents, are clearly distinct from the tholeiitic and Mand C- series rocks but overlap with the calc-alkaline field
and extend it towards more more primitive compositions.
Trace element ratios, such as Nb/Yb and Ba/La (not
shown) also distinguish low-K tholeiites and Soufriere
eruptives. We suggest, therefore, that to distinguish them
from the tholeiitic rocks of the northern islands, the rocks
from Soufriere are best referred to as having calc-alkaline
character, and that some may represent the mafic endmember of the calc-alkaline magma series in the Lesser
Antilles.
The Soufriere rocks are close compositionally to a suite
of lavas in the neighbouring island of Bequia named the
isotopically homogeneous suite (IHS) by Smith et al.
(1996), who pointed out that the similarities suggest that
the two suites were derived from similar sources and
have similar evolutionary histories.
Magmatic affinity
Figure 11 is a plot of MgO vs degree of silica saturation.
The Soufriere St Vincent rocks form an array stretching
from silica saturation (Q ~0) to silica oversaturation (Q
~20), overlapping at the higher-Q end with the calcalkaline and low-K tholeiite fields. The slope of the array,
and the slightly nepheline-normative character of STV
301, seem to indicate that more primitive Soufriere
magmas, if they existed, would have been nephelinenormative, a feature which they would share with the
Major and trace element variations
Major and selected trace element variations in Soufriere
rocks are displayed on MgO plots in Fig. 14. Six Yellow
Tuff Formation samples, five of which were separated
from palpably mixed rocks, define a linear, mixing, trend.
The end-members are a basalt with MgO = 10 wt %
and a silicic andesite (MgO = 2 wt %), which represents
the most evolved rock erupted from the Soufriere. In the
other rocks, SiO2, Na2O, K2O, P2O5, Ba, Rb, Sr, Th,
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Fig. 10. SiO2–CaO plot for glass (melt) inclusions trapped in phenocrysts (host phenocrysts shown by different symbols). Also shown is the field
of whole-rock analyses (stippled area). The comparability of inclusion and rock data in products of the 1979 eruption was noted by Graham &
Thirlwall (1981), Devine & Sigurdsson (1983) and Bardintzeff (1984, 1992). Data sources: this study (Table A6); additional analyses of olivinehosted melt inclusions with SiO2 <55 wt % from Graham & Thirlwall (1981) and Devine & Sigurdsson (1983).
Fig. 11. Degree of silica saturation (CIPW normative) plotted against MgO for Soufriere rocks and the main Lesser Antilles suites. Q is
normative quartz plus quartz in hypersthene. Lesser Antilles data sources: Wills (1974), Hawkesworth et al. (1979), Graham (1980), Graham &
Thirlwall (1981), Davidson (1984, 1985, 1987), Thirlwall & Graham (1984), White & Patchett (1984), White & Dupré (1986), Davidson &
Harmon (1989), Thirlwall et al. (1994, 1996), Smith et al. (1996), Turner et al. (1996), and R. Macdonald (unpublished data, 1997).
U, Zr, Hf and rare earth elements (REE) increase, and
CaO, Fe2O3∗, Co, Cr, Ni, Sc and V decrease, with
decreasing MgO. The distribution of TiO2, and possibly
also Cs and Nb, is more complex and seems to show a
decrease to 6 wt % MgO and then an increase with
further decrease in MgO. A rather similar TiO2 pattern
seems to be shown by the IHS on Bequia (Smith et al.,
1996, fig. 6). V abundances in the IHS, however, increase
with increasing differentiation.
Chondrite-normalized REE patterns (Fig. 15) are gen-
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HEATH et al.
MAGMAGENESIS AT SOUFRIERE VOLCANO
Fig. 12. (a) SiO2–FeO∗/MgO classification diagram (Miyashiro, 1974). The slope of the Soufriere data does not indicate a clear affinity with
either the tholeiitic or calc-alkaline series. (b) K2O–SiO2 classification diagram. The dotted lines enclose the field boundaries between (low-K)
tholeiitic and calc-alkaline rocks collated by Rickwood (1989). (c) AFM plot [total alkalis (Na2O + K2O)–total iron (FeO + Fe2O3)–MgO] of
Soufriere eruptive rocks. Field boundary is from Irvine & Baragar (1971).
erally flat to slightly light REE (LREE) enriched ([La/
Yb]N 0·9–2·2), with weak Eu anomalies which range
from slightly positive in the most magnesian basalts (Eu/
Eu∗ Ζ1·07) to slightly negative in the basaltic andesites
([0·96) and andesite (0·89). With one exception, the
rocks show a Ce anomaly, Ce/Ce∗ ranging from 0·87
to 0·98. Flat patterns like those from Soufriere have
previously been recorded from calc-alkaline basalts and
basaltic andesites from St Lucia and Guadeloupe, the lowK tholeiites of St Kitts and certain silica-undersaturated
basalts of Grenada (White & Dupré, 1986; Thirlwall et
al., 1994).
With decreasing MgO content of Soufriere whole rocks,
LREE/MREE (medium REE) (e.g. La/Sm) and LREE/
HREE (heavy REE) (La/Yb) ratios decrease to rocks
with MgO around 3 wt %, with a sharp increase into
the andesite STV 376(L) (Fig. 16). This is unusual behaviour in a suite of arc rocks, where LREE enrichment
with increasing differentiation is the norm, as for example,
in the IHS of Bequia (Smith et al., 1996). There is,
however, apparent LREE depletion in more differentiated members of the isotopically diverse suites (IDS)
of lavas on Bequia.
Figure 17 presents chondrite-normalized element
abundance plots for three basalts with MgO >7 wt %.
The patterns are characteristic of arc magmas, namely,
high LILE/HFSE (large ion lithophile elements/high
field strength elements) and LREE/HFSE ratios, Nb
depletion relative to La, and relatively high Ba/La and
Sr/Nd ratios. The LREE sections of the patterns are
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Fig. 13. 87Sr/86Sr vs MgO content for Lesser Antilles volcanic rocks subdivided on the basis of magmatic affinity. Lesser Antilles data sources
as for Fig. 11. The field of Atlantic N-type MORB is from Saunders et al. (1988).
relatively flat, their abundance varying from 14 to 18
times chondritic. The HFSE and HREE sections are also
rather flat, with 8–12 times enrichment over chondrites.
The rocks are enriched in the LILE (Rb, Th, K and Ba),
but not Sr, compared with the LREE. Average MORB
is also plotted in Fig. 17. As is typical of arc magmas,
the Soufriere rocks show LILE and LREE enrichment
and HFSE depletion relative to MORB.
Isotopic data
Figures 18 and 19 show Sr–Nd–Pb isotopic variations in
Lesser Antilles rocks, including our new data from Soufriere St Vincent, and also fields for Atlantic Ocean
sediments and typical MORB. 87Sr/86Sr ratios (0·703752–
0·704407) of Soufriere rocks are somewhat higher than
those of MORB, whereas 143Nd/144Nd ratios (0·512976–
0·512843) are slightly lower. The new Pb data extend
the range established by previous workers (White &
Dupré, 1986; M. F. Thirlwall, unpublished data figured in
Smith et al., 1996) to slightly less radiogenic compositions
(206Pb/204Pb = 18·644–19·328). On Pb–Pb plots (Fig. 19),
the Soufriere data (and the closely similar IHS of Bequia,
not shown) form linear trends which extrapolate to intersect the fields of MORB and locally subducting sediments, as determined from Deep Sea Drilling Project
(DSDP) Hole 543. This strongly suggests that at least
two components have been involved in their petrogenesis—depleted mantle (MORB) and a subduction
component at least partly derived from the sediments
(White & Dupré, 1986; Smith et al., 1996).
The relative homogeneity of the Soufriere and Bequia
IHS rocks compared with calc-alkaline rocks of the
central islands (St Lucia, Martinique, Dominica) is mainly
a function of MgO content; isotopic diversity within
series is restricted to rocks with MgO <5 wt % (Fig. 13).
However, at given MgO, Soufriere St Vincent and other
calc-alkaline rocks of the Lesser Antilles arc typically
have higher 87Sr/86Sr, 206Pb/204Pb and 207Pb/204Pb ratios,
and lower 143Nd/144Nd than the tholeiitic lavas of the
northern islands. The M- and C-series lavas of Grenada
typically have higher 87Sr/86Sr and lower 143Nd/144Nd at
given MgO than the calc-alkaline rocks. As these features
characterize even the most primitive rocks, they are likely
to indicate isotopically distinct mantle sources.
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MAGMAGENESIS AT SOUFRIERE VOLCANO
Fig. 14. (a).
Magmatic lineages
As a result of scatter within the rocks of each formation,
it is difficult to distinguish between the different formations at Soufriere on the basis of major or trace element
compositions. There are, though, some minor differences.
For example, the Pre-Somma lavas generally have slightly
higher K2O abundances at given MgO than the Crater
Lavas and rocks of the Pyroclastic Formation (Fig. 14a).
On a more detailed scale, Graham & Thirlwall (1981)
noted that the 1971 lava is depleted in Al, Ca and Sc,
enriched by ~10% in Na, K, Rb, Zr and Zn, and higher
in Cr, compared with the 1979 lava. They suggested that
the lavas represent two entirely different magma batches.
Magnesian basalt STV 301 has higher Ti, K, P, La, Nb,
Th and U, and lower Ni, abundances, and lower Rb/K
ratios, than Pre-Somma rocks of similar MgO content.
In contrast, STV 315 (MgO = 7·8 wt %) has low
abundances of La, Sr and Th compared with rocks
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Fig. 14. (a) Major element and (b) selected trace element abundances in Soufriere rocks, plotted against MgO content. The continuous lines
join the samples whose major element compositions are used in least-squares fractionation modelling.
of similar MgO content. As there are no systematic
correlations between trace element ratios, or between
them and any isotopic ratio, we infer that the mantle
sources were compositionally heterogeneous on a small
scale.
Thus, although we discuss the Soufriere rocks as describing a cogenetic suite, in reality they probably rep-
resent many different liquid lines of descent, a point
already made for the recent pyroclastic deposits of Mt
Pelée, Martinique (Smith & Roobol, 1990) and the Mseries rocks of Grenada (Devine, 1995; Thirlwall et al.,
1996).
A few generalizations may be made concerning the
distribution of rock types in time. The most primitive
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MAGMAGENESIS AT SOUFRIERE VOLCANO
Fig. 15. Chondrite-normalized REE abundances in selected Soufriere
basalts. Normalizing factors from Nakamura (1974), except for Tb
(Wakita et al., 1971).
Fig. 17. Primordial mantle-normalized trace element plots for selected
Soufriere basalts. Normalizing values are from McDonough et al. (1992),
except P (92 ppm; Sun, 1980). MORB data are from Pearce (1983),
except La, Nd and Tb, from Saunders & Tarney (1984).
Fig. 16. [La/Sm]N and [La/Yb]N against MgO for Soufriere rocks. It should be noted that both ratios decrease overall with increasing
fractionation. The dashed line in the [La/Yb]N plot gives the results of modelling Rayleigh fractionation, as follows: parent STV 358 to daughter
STV 351; phenocryst assemblage is 13·8% ol + 21·8% cpx + 37·6% pl + 1·3% Fe–Ti oxides. Partition coefficients (from Dostal et al., 1983,
table 5): ol (La 0, Yb 0), cpx (0·13, 0·60), pl (0·10, 0·02), Fe–Ti oxides (0, 0). The model predicts LREE enrichment, rather than the observed
depletion, relative to HREE. Inclusion of amphibole into the fractionating assemblage (La 0·25, Yb 1·20) increases the LREE enrichment.
lavas (SiO2 <50 wt %) are restricted to the Pre-Somma
Formation and the average composition of Pre-Somma
rocks is accordingly less SiO2 rich than younger rocks.
The widest compositional range, basalt to andesite, occurs
in the Yellow Tuff Formation and we can infer that the
magma reservoirs were more mature by that stage. More
recent activity (the Crater Lavas and Pyroclastic Formation) has been dominated by basaltic andesites but
eruption of basalt during the compositionally zoned 1902
eruption suggests that basalt magma has been available
at depth in the system, although normally unable to
erupt through more evolved overlying magmas.
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Fig. 18. 87Sr/86Sr–143Nd/144Nd plot for Soufriere eruptive rocks and xenoliths and for other Lesser Antilles suites and MORB. Data sources as
in Figs 11 and 13.
MELT INCLUSIONS
Graham & Thirlwall (1981), Devine & Sigurdsson (1983)
and Bardintzeff (1992) provided compositional data on
melt inclusions in products of the 1979 eruption. We add
here (Table A6) electron microprobe data for a further
eight rocks, including representatives of all the major
stratigraphic units. The following discussion incorporates
material from the earlier papers.
Melt inclusions are found in all the main phenocryst
phases. Compositions range from basalt (47 wt % SiO2)
to high-silica rhyolite (77 wt % SiO2). A considerable
part of the range may be found in the products of single
eruptions. Thus, whole-rock analyses of the products of
the 1979 eruption show very little major element variation
(Graham & Thirlwall, 1981, table 1), whereas melt inclusions in phenocrysts range in SiO2 content from 48·6
to 62·0 wt % (Devine & Sigurdsson, 1983, table V).
The melt inclusions extend to high-silica compositions
unknown among the eruptive products of Soufriere and
relatively uncommon in the Lesser Antilles arc. It is also
important to note that inclusions with much higher silica
contents than the host rocks occur in every sample
analysed, e.g. inclusions of rhyolitic composition occur
in basalt STV 315. It is possible that melt compositions
in such cases have been significantly modified by postentrapment crystallization. This effect is likely to have
been small at Soufriere, however. Graham & Thirlwall
(1981) noted, for example, that the melt inclusions in
olivine do not show the distribution of compositions along
an olivine control line expected from further crystallization of olivine after inclusion. Similar comments
can be made for the absence of plagioclase control
among inclusions in plagioclase phenocrysts (see Devine
& Sigurdsson, 1983, fig. 2). From a detailed analysis of
differences in inclusion composition in different phenocryst hosts, Devine & Sigurdsson (1983) concluded
that, on average, no more than 8 wt % of any inclusion
could have precipitated out of the host mineral during
quenching. These observations, plus the absence of optical and electron probe evidence for crystallization of the
inclusion rinds, lead us to conclude that post-entrapment
crystallization has not significantly altered the compositions of the melt inclusions.
The following observations suggest, on the other hand,
that the inclusions represent trapped liquids which were
cognate to the host phenocrysts:
(1) There is an overall positive correlation between the
SiO2 contents of inclusions and host rocks (Tables A2
and A6).
(2) There is also a relationship between inclusion
composition and the order of crystallization of phenocryst
phases. Thus, olivine phenocrysts tend to have inclusions
which are more magnesian than those in plagioclase,
whereas those in orthopyroxene are almost exclusively
more evolved (Devine & Sigurdsson, 1983; Bardintzeff,
1992).
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MAGMAGENESIS AT SOUFRIERE VOLCANO
inclusions in rocks of the Fuego 1974 eruption, Guatemala. Devine (1995) reported, however, Cl abundances
between 0·34 and 0·38 wt % in melt inclusions in Mseries lavas from Grenada. Relatively high Cl contents
may be a feature of Lesser Antilles magmas.
PRE-ERUPTIVE WATER CONTENTS
OF THE MAGMAS
In this section, we estimate the pre-eruptive water contents of Soufriere magmas using three techniques: (1)
the ‘difference’ method from published electron probe
analyses; (2) comparison with phase equilibrium experimental results; (3) clinopyroxene compositions.
The difference method
Fig. 19. Plots of 206Pb/204Pb vs (a) 207Pb/204Pb and (b) 208Pb/204Pb in
Soufriere rocks (plotted symbols represent different geological formations, as in Fig. 14). Also shown are fields for MORB and locally
subducting sediments and the Northern Hemisphere Reference Line
(NHRL; all from Smith et al., 1996).
(3) Major element variations in the inclusions mimic
those in the whole rocks (Fig. 10).
The matrix compositions of Soufriere basaltic andesites
and andesites are invariably silicic (Graham & Thirlwall,
1981; Devine & Sigurdsson, 1983; Bardintzeff, 1992) and
in the only rocks for which melt inclusion and matrix
glass data are available, those of the 1979 eruption,
the melt inclusions extend to much more magnesian
compositions than the matrix glass (e.g. Devine & Sigurdsson, 1983, table V). We suggest, therefore, that
growing phenocrysts incorporated increasingly evolved
residual melts.
In any given Soufriere sample, the highest Cl abundances are found in melt inclusions in clinopyroxene
phenocrysts. If we take these to be closest to magmatic
Cl abundances, they indicate contents mainly between
0·2 and 0·3 wt % and up to 0·6 wt %. Such values seem
to be rather high for arc magmas; for example, Perfit et
al. (1980) reported a range of 460–2000 ppm Cl in calcalkaline magmas, and Sisson & Layne (1993) found, by
ion microprobe, Cl abundances <0·16 wt % in melt
The difference between 100% and the analytical total
for melt inclusion compositions derived from electron
probe analysis may be used as an indicator of the volatile
content of the inclusions (Anderson, 1979). Devine &
Sigurdsson (1983) and Bardintzeff (1992) have applied
the technique to products of the 1979 eruption of Soufriere, the former workers giving a detailed analysis of the
potential problems of interpretation. They concluded
that the difference method can give reliable estimates of
volatile content (inferred to be dominated by water), a
conclusion supported by Sisson & Layne (1993), who
showed that summation deficits in inclusions from various
subduction-related volcanoes agree with the water contents determined by ion microprobe with an average
error of ±0·65 wt %.
Devine & Sigurdsson (1983) reported that the volatile
content of olivine-hosted inclusions in the 1979 Soufriere
rocks ranged from 2 to 8 wt %, averaging 2·9 wt %. In
contrast, the volatile content of inclusions in plagioclase
and pyroxenes was probably <2 wt %, suggesting that
the more evolved magmas had degassed before eruption.
Bardintzeff (1992) also found a negative, but very
scattered, correlation between inferred water content and
SiO2 content of the glasses in 1979 products. Rather than
using data from inclusions in phenocrysts as indicators of
water contents in the basaltic magmas, he chose instead
data for mafic interstitial glasses in cumulate blocks,
giving an average of 2·5 wt %, comparable with the
Devine & Sigurdsson estimate. For more evolved rocks,
Bardintzeff (1992) chose, somewhat arbitrarily, a range
of 3–5 wt %, with an average of 4 wt %, i.e. water
contents increased with differentiation.
Phase equilibrium studies
We now compare magma compositions with experimentally determined phase equilibria in the system
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olivine (Ol)–high-Ca pyroxene (Cpx)–plagioclase (Pl)–
silica (Qz). This analysis is based mainly on the data
and discussion of Sisson & Grove (1993a, 1993b), who
determined the phase relations of a series of natural
aphyric high-alumina basalts and intrusive equivalents
through melting experiments at 2 kbar, under watersaturated conditions and with f (O2) buffered at NNO
(nickel–nickel oxide), i.e. at somewhat less oxidizing conditions than prevailed at Soufriere.
Figure 20 is the pseudoternary Ol–Cpx–Pl, projected
from Qz–magnetite (Mt)–orthoclase (Or)–apatite (Ap). It
shows the hydrous 2 kbar multiple saturation boundary
for ol + cpx + pl, deduced from experiments on natural
basalts and projected back towards the piercing point in
the forsterite–diopside–anorthite system. Sisson & Grove
(1993a) noted that the compositions of high-alumina
basalts, basaltic andesites and andesites from the Aleutians
and Fuego volcano in Guatemala plot close to the watersaturated multiple saturation boundary, and suggested
that the close correspondence indicates that the magmas
had pre-eruptive water contents comparable with those
of the experimental liquids (up to 6 wt %).
Basalts and basaltic andesites from Soufriere St Vincent
are also plotted in Fig. 20. A first interpretation would
be that the series evolved by ol + cpx crystallization to
the hydrous 2 kbar multiple saturation boundary, where
plagioclase also began to crystallize. Such an interpretation is not, however, totally supported by petrographic evidence. Almost all the Soufriere suite is
plagioclase-phyric, not just those plotting on the multiple
saturation boundary.
Similar discrepancies are seen in other projections
in the same system. Figure 21 is a projection from
Pl–Mt–Or–Ap onto Ol–Cpx–Qz. Liquids saturated with
ol + high-Ca cpx + pl (± spinel) define a multiple saturation boundary which projects from the silica-poor
side of the Ol–Cpx sideline to the Ol–Qz sideline. Sisson
& Grove (1993a) noted that, as in the last projection,
there is a close correspondence between the position of
the boundary and the fields of high-alumina basalts and
andesites from the Aleutians and Fuego volcano, and
again took this as evidence that the natural magmas had
evolved under water-rich conditions. Figure 21 also shows
the multiple saturation boundary defined by liquids saturated with ol + pl + hb (± mt), produced in melting
experiments on hornblende-bearing intrusive rocks (Sisson & Grove, 1993a). The boundary in this case is a
reaction boundary along which olivine and liquid react
to produce hornblende. When all the olivine is consumed,
the liquids follow a pl + hb + mt coprecipitation trend
towards the Qz apex.
Whole-rock data from Soufriere are projected onto the
Ol–Cpx–Qz plane in Fig. 21. They start on the silicapoor side of the Ol–Cpx sideline, cross the multiple
saturation boundary and then trend towards the Qz
NUMBER 10
OCTOBER 1998
apex. The more evolved part of the trend thus appears
similar to the hb + pl + mt control line. There are at
least two problems with this interpretation: (1) amphibole
is not a phenocryst phase in less evolved rocks, as would
be required if the liquid trend were controlled by pl + hb
crystallization; and (2) Sisson & Grove (1993a) confirmed
the results of Kushiro (1969) that high contents of dissolved water destabilize low-Ca pyroxene relative to
olivine, and they suggested that low-Ca pyroxene will
not crystallize from water-rich basaltic or basaltic andesite
melts. At Soufriere, however, orthopyroxene starts crystallizing from basaltic andesite liquids whose projected
compositions lie on the multiple saturation boundary
ol + cpx + pl + liquid. Crystallization of the assemblage
ol + pl + cpx + opx + Ti-mag then drove liquids down
towards the quartz apex. Thus, although the broad
spread of analyses around the 2 kbar multiple saturation
boundary is consistent with high water contents in the
Soufriere magmas, the early crystallization of orthopyroxene points to water-poor conditions.
This lack of correspondence between projected composition and phenocryst assemblage has been noted by
Feeley & Davidson (1994), who, in a study of calc-alkaline
lavas and magmatic inclusions of the Volcán Ollagüe in
the central Andes, found that basaltic andesite lavas
scatter around 2 kbar ol + hb + pl and ol + cpx + pl
saturation boundaries in the Ol–Cpx–Qz pseudoternary,
despite the facts that they are not plagioclase or amphibole
saturated and that olivine phenocrysts are rare or absent
in the magmatic inclusions. Those workers suggested that
the correspondence is coincidental, the trend probably
reflecting crustal assimilation or mixing with or without
fractional crystallization.
We now use an alternative set of phase relationships.
Figure 22 shows pseudoternary phase diagrams in the
system Ol–Di–Pl–Mt–SiOr, where SiOr is a ‘magmaphile’ component consisting of combined normative
quartz and orthoclase. Phase relationships were determined from a suite of high-alumina basalts and andesites from Atka island in the Aleutians (Baker & Eggler,
1987). Experimental conditions varied from anhydrous,
at pressures from 1 atm to 8 kbar, to hydrous (2 wt %
H2O in the melt) at 2 and 5 kbar.
Soufriere rocks are plotted using the algorithm of
Baker & Eggler (1983), with Fe2O3/FeO = 0·22. In the
projection from Di–Mt (Fig. 22a), presented by Baker &
Eggler (1987) as an effective geohygrometer, the basaltic
andesites plot close to the 2–5 kbar hydrous liquid lines
of multiple saturation (LLMS) but the basalts clearly do
not lie on any reasonable extrapolation of the trends.
The phase relationships seem to point to relatively dry
basaltic magmas at Soufriere. In the Pl–Mt projection
(Fig. 22b), which is more useful as a geobarometer (Baker
& Eggler, 1987), the basaltic andesites and andesites of
Soufriere St Vincent lie close to the 5 kbar hydrous
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MAGMAGENESIS AT SOUFRIERE VOLCANO
Fig. 20. Pseudo-ternary projection, from Qz into the plane Ol–Pl–Cpx, of multiple saturation boundaries representing liquids saturated with
ol + high-Ca cpx + pl, showing how projected positions change with increasing H2O. The projection scheme is that of Tormey et al. (1987),
including corrections by Grove (1993), and the boundaries were constructed from experiments on natural liquids (Sisson & Grove, 1993a, and
references therein). The dot–dashed lines enclose the field of 1 atm anhydrous boundaries, the dotted line is for 2 kbar water-undersaturated
conditions (pH2O = 0·7 ptotal) and the continuous line is the boundary for 2 kbar water-saturated conditions, projected back towards the
piercing point in the forsterite–diopside–anorthite system. Increasing H2O therefore shifts the multiple saturation boundary away from the Cpx
apex, and to a lesser extent away from the Ol apex. Also shown are rocks from Soufriere, symbols representing phenocryst assemblages.
LLMS for ol + pl + aug and pl + aug + opx. The distribution is consistent with an origin from basaltic parents
by fractionation of ol + pl + aug, to produce derivative
melts with ~2 wt % H2O. The array of Soufriere basalts
is again more consistent with crystallization under waterpoor conditions, at pressures exceeding 8 kbar.
Clinopyroxene compositions
Experiments by Gaetani et al. (1993) on the crystallization
of basalts containing up to 6 wt % dissolved water have
provided some evidence on the effects of water on the
compositions of high-Ca pyroxenes. The effect of increasing pressure up to 15 kbar in anhydrous systems
is to decrease the Wo content of clinopyroxene while
increasing (CaTs + CrTs), the Al- and Cr-rich components (Fig. 23). In the experiments of Gaetani et al.,
conducted at 2 kbar, the effect of 6 wt % dissolved water
was to increase the Wo content of the clinopyroxenes at
nearly constant CaTs + CrTs. The fields of clinopyroxenes in ol- and cpx-phyric lavas from Soufriere and
from selected arcs are shown in Fig. 23. The similarity
to the clinopyroxenes produced experimentally at 2 kbar
led Gaetani et al. (1993) to suggest that the arc lavas
had evolved through hydrous crystallization at crustal
pressures. A similar conclusion can be made for the
Soufriere rocks, despite the fact that they probably equilibrated at higher pressures.
In summary, the various methods for estimating magmatic water contents have given conflicting results. Clinopyroxene compositions are consistent with water contents
exceeding 3 wt %, whereas the early crystallization of
plagioclase and orthopyroxene is more consistent with
relatively anhydrous basaltic magmas generating hydrous, but water-undersaturated, andesitic magmas. The
absence of amphibole phenocrysts may also point to
water-undersaturated conditions. This is unlikely to be
due solely to a compositional effect, as many Soufriere
basaltic andesites contain >3 wt % Na2O, the level at
which amphibole should be stabilized, given high enough
water contents (Cawthorn & O’Hara, 1976; Sisson &
Grove, 1993a). Given these uncertainties, it is still unclear
whether magmatic evolution at Soufriere was ac-
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Fig. 21. Pseudo-ternary projection, from Pl into the plane Ol–Cpx–Qz, of various multiple saturation boundaries (Sisson & Grove, 1993a). (1)
represents liquids saturated with olivine, high-Ca clinopyroxene and plagioclase under 2 kbar, water-saturated conditions, produced in 2 kbar
water-saturated experiments on natural high-Al2O3 basalts. The large star is a distributary reaction boundary from which two hornblendesaturated boundaries exit. (2) Liquids saturated with olivine, hornblende and plagioclase define a reaction boundary along which olivine and
liquid react to form hornblende. Liquids were produced in 2 kbar, water-saturated experiments on a hornblende gabbro. (3) Hornblende–plagioclase
control generates corundum-normative liquids (off the diagram, in a direction indicated by the arrow). (4) The 1 atm pressure ol + px + pl
multiple saturation boundary. The hachured area shows the stability field of orthopyroxene at 1 atm. Different phenocryst assemblages in
Soufriere rocks are shown by different symbols. The projection scheme is that of Tormey et al. (1987) and Grove (1993), using mineral components
in oxygen units.
companied by continuous (or episodic) degassing, or by
increasing water contents in residual melts.
Magmatic evolution
Many previous studies of the evolution of primary
magmas in Lesser Antilles volcanic suites have recognized
the polybaric nature of the fractionation processes and
the role on many islands of crustal assimilation (Thirlwall
& Graham, 1984; White & Patchett, 1984; Davidson,
1985, 1986, 1987; White & Dupré, 1986; Davidson &
Harmon, 1989; Smith & Roobol, 1990; Bardintzeff, 1992;
Thirlwall et al., 1994, 1996; Smith et al., 1996; Turner et
al., 1996). Below, we comment on the nature of the
primary magmas, assess the degree of crustal contamination in the Soufriere rocks and then attempt to
model magma evolution by fractional crystallization.
First, we should comment on the role of magma mixing
in the evolution of the Soufriere suite. As noted in the
Petrography section, macroscopically mixed rocks have
been found at a few localities on Soufriere and this must
raise the possibility that other members of the suite
represent more thoroughly hybridized materials. The
spread of analyses in MgO variation diagrams (Fig. 14)
would allow for this possibility. We note, however, the
following points:
(1) Reverse zoning in phenocrysts is, however, rare to
absent at Soufriere, except as part of oscillatory zoning
in plagioclase. This is not consistent with incorporation
into the magmas of large volumes of melt of different
temperature.
(2) The non-linear nature of compositional variation
within the suite (e.g. MgO–Na2O and MgO–CaO,
Fig. 14a) is more consistent with evolution by fractional
crystallization.
(3) There is an overall correlation between phenocryst
and whole-rock compositions.
Although, therefore, there may have been a minor role
for mixing, or for heat transfer from new influxes of
basalt lower into the magma reservoirs, we suggest that
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MAGMAGENESIS AT SOUFRIERE VOLCANO
Fig. 22. (a) The Ol–Pl–SiOr pseudoternary projected from Di–Mt, and (b) the Ol–Di–SiOr pseudoternary, projected from Pl–Mt, with phase
relationships deduced from melting experiments on rocks from Atka island (Baker & Eggler, 1987). The dashed lines are the 1 atm
multiple saturation boundaries ol + aug + pl + melt. The continuous lines are the 8 kbar anhydrous boundaries ol + pl + pig + melt,
ol + pl + aug + melt and pl + pig + aug + melt. The dotted lines are the 2 kbar, water-undersaturated (2 wt % H2O) multiple saturation
boundaries ol + pl + aug + melt, ol + aug + opx + melt, which meet at an olivine–orthopyroxene reaction point. The dot–dashed lines show
the same boundaries at 5 kbar, water undersaturated. The projection scheme is that of Baker & Eggler (1983). Soufriere rocks are plotted using
different symbols to represent the main geological formations.
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Fig. 23. Clinopyroxene ternary diagram, with compositions recalculated into end-members Wo (CaSiO3), En + Fs [(Mg,Fe)SiO3] and
CaTs + CrTs [Ca(Al,Cr)SiAlO6]. The broad arrow gives the trend of pyroxene compositions produced in anhydrous melting experiments
between 1 atm and 15 kbar. The dashed-line field encloses the compositions of pyroxenes produced in 2 kbar, water-saturated melting experiments
on (ol + cpx)-phyric basalts from the Lau basin and natural clinopyroxene phenocrysts from (ol + cpx)-phyric arc lavas. All after Gaetani et al.
(1993). Also shown are clinopyroxene data from (ol + cpx ± plag)-phyric Soufriere basalts and basaltic andesites.
magma mixing has not been a major differentiation
process at Soufriere.
Primary magma compositions
Identification of primary magmas is critically dependent
on the composition chosen for the mantle source rocks,
in particular the mg-number of the olivine. Some simple
whole-rock compositional criteria are widely used, including mg-number >70, FeO∗/MgO <1, Ni >200 ppm
and Cr >400 ppm (Tatsumi & Eggins, 1995). These
criteria are met by the more magnesian lavas of the PreSomma group (e.g. STV 301, 310 and 358) and we have
already shown that olivine phenocrysts in STV 301 were
in (near-)equilibrium with the whole-rock composition.
These rocks may therefore be close to primary magma
compositions. It has already been inferred that the parental rocks of St Vincent suites and of the IHS of Bequia
contained >12 wt % MgO (Smith et al., 1996), although
normative compositions (Fig. 11) may indicate that the
melts were nepheline-normative. In comparison,
Thirlwall et al. (1996) argued that all Grenada magmas
were ultimately derived by fractional crystallization from
undersaturated picrites with ~16 wt % MgO.
Degrees of partial melting
Geochemical studies have established that there have
been three components in the mantle sources of Lesser
Antilles magmas: (1) the mantle wedge, which is similar
to, or slightly enriched in HFSE relative to, the N-MORB
source; (2) a fluid phase added to the wedge from the
subducting slab, with contributions from both sediments
and altered basaltic crust; (3) a component formed by
partial melting of subducted sediments (Thirlwall & Graham, 1984; White & Patchett, 1984; Davidson, 1986,
1987; White & Dupré, 1986; Davidson & Harmon, 1989;
Thirlwall et al., 1994, 1996; Smith et al., 1996; Turner et
al., 1996; Hawkesworth et al., 1997). Although it is
accepted that the proportions of each component vary
along the arc, there is as yet no consensus on how the
variation relates to individual islands.
There are major difficulties associated with estimating
degrees of partial melting of mantle sources, related, inter
alia, to uncertainties about mantle composition and the
effects of water content on phase equilibria. For example,
Stolper & Newman (1994) have suggested, in a study of
glasses from the Mariana trough, that a 0·2 wt % increase
in water content in the source of the melts caused the
degree of melting to rise by some 12 wt %.
In the case of the Lesser Antilles, La/Yb has been
used qualitatively by Thirlwall et al. (1994) as a measure
of degree of partial melting, higher ratios signifying
smaller degree melts. Those workers suggested that St
Vincent rocks, on the basis of relatively low La/Yb ratios,
represent fairly high-degree melt fractions, which they
saw as consistent with their transitional tholeiitic–calcalkaline character.
Pearce & Parkinson (1993) have used an Nb–Yb plot
to study the melting process in arcs, and argued that the
element–element plot allowed them to separate source
depletion from degree of melting. To minimize the effects
of fractional crystallization, which may be significant for
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MAGMAGENESIS AT SOUFRIERE VOLCANO
HFSE even in magmas with MgO between 10 and 6 wt
% (Thirlwall et al., 1994), Pearce & Parkinson (1993)
recommended normalizing Nb and Y values to an MgO
content of 9 wt %. Relevant values for Soufriere rocks
are Nb 3 ppm and Yb 2·2 ppm, which suggests a degree
of partial melting of 15%. Given the rather dubious
assumption in such plots that a given MgO content
signals equal amounts of fractionation of the primary
magmas (Thirlwall et al., 1994), the value of 15% can
be taken as no more than an indicator of the degree of
melting.
An important result from the study by Thirlwall et al.
of HFSE anomalies is that there is no evidence for residual
minor phases in the source mantle of the Soufriere St
Vincent magmas. For example, near-constant Ti/Eu,
Zr/Sm and Nb/La ratios during partial melting, as
inferred from La/Yb ratios, preclude a role for residual
amphibole, as amphibole should fractionate HFSE from
REE. If such phases existed in the mantle sources, they
disappeared during the melting process.
Crustal contamination
Various geochemical criteria have been used to assess
the effects of crustal contamination during magma evolution in the Lesser Antilles. These include increases, with
advancing differentiation, in 87Sr/86Sr and 206Pb/204Pb,
decreases in 143Nd/144Nd, and increases in Rb/Ba, Rb/
K, K/La, La/Yb, Pb/Sr and U/La (Thirlwall & Graham, 1984; White & Dupré, 1986; Davidson, 1987;
Davidson & Harmon, 1989; Smith et al., 1996; Thirlwall
et al., 1996).
Sr and Nd isotopic ratios in Soufriere rocks have only
a limited range, are close to MORB in composition and
show no systematic change with degree of fractionation
(Figs 13 and 18). This suggests that the magmas evolved
by fractional crystallization in a closed system. It is
also possible that open system assimilation–fractional
crystallization (AFC) occurred where the magmas and
assimilant(s) had similar isotopic characteristics. In this
regard, we note that some of the xenoliths analysed
during this study are isotopically indistinguishable from
the lavas (Table A5). However, incompatible trace element ratios in evolved rocks show the same range as the
most magnesian basalts, as demonstrated, for example,
by Rb/K–SiO2 relationships (Fig. 24; after Davidson,
1987), suggesting that the range has been derived from
the mantle source(s). The spread in the data means that
we cannot preclude a contamination component, but we
equally can conclude that there is no compelling evidence
that the Soufriere magmas experienced significant crustal
contamination.
Modelling fractionation trends
Major element variations (except P) within the suite have
been modelled using the Mix ’n’ Mac petrological mixing
program by D. R. Mason, which is based on the leastsquares method of Bryan et al. (1969). Sums of squared
residuals (Rr 2) should be <0·2 for an acceptable solution.
Modelling was done in three steps: magnesian basalt
STV 310 to basalt STV 334 to basaltic andesite STV
79-70 to andesite STV 376(L). The whole-rock compositions are joined by continuous lines in Fig. 14. The
mineral data used in the modelling were selected to
reflect typical equilibrium compositions with respect to
the parent magma at each stage. Input Fe2O3/FeO
ratios varied with the silica content of the whole rock
(Middlemost, 1989).
The best solutions (i.e. those with low Rr 2 and reasonable degrees of crystallization) were achieved by fractionating 14% ol + 8% cpx + 14% plag + 1% Ti-mag
± 0·1% Cr-sp from basalt STV 310, then 8% ol + 10%
cpx + 9% plag + 0·5% Ti-mag from basalt STV 334,
followed by 7% cpx + 26% pl + 4% Ti-mag + 7% opx
from basaltic andesite STV 79-70. These are the observed
phenocryst assemblages, and the relative proportions of
fractionating phases are similar to those found in the
rocks (Table A1). Given the range of phenocryst compositions and proportions in the rocks and the fact that
the suite comprises many liquid lines of descent, such
models cannot be considered robust [a point made by
Devine (1995) in modelling the evolution of Grenada
rocks]. Nevertheless, the results are at least consistent
with evolution of the Soufriere suite by closed-system
fractionation of the observed phenocryst phases, the most
evolved rock representing some 76 wt % crystallization
of the parental basalt.
We have not attempted to model the derivation of the
rhyolitic melt inclusions from andesitic compositions.
However, Devine & Sigurdsson (1983) found that the
interval 60–75 wt % SiO2 could be satisfactorily modelled
using assemblages close to 5% ol + 13% cpx + 74%
pl + 6% Ti-mag + 2% opx. Because of the reaction
relationship between them, it is unlikely that olivine
and orthopyroxene coprecipitated in the fractionating
assemblage. The important observation, though, is that
compositional variation in the suite can apparently be
explained by removal or addition of the observed phenocryst phases.
Apart from some aspects of REE distribution, we have
not attempted to model trace element variations in the
Soufriere rocks because of the shortage of partition coefficient data for the phenocrysts. Trace element distributions are, however, qualitatively consistent with
fractional crystallization of the observed phases. The
rapid decrease of Ni and Cr, and the slower decrease
in Co and Sc, with slowly increasing Zr is typical of
fractionation of olivine and clinopyroxene, a point made
for the IHS of Bequia by Smith et al. (1996). The
systematic increases in ITE from most to least magnesian
rocks are also consistent with an evolution by fractional
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Fig. 24. Rb/K vs SiO2 plot, after Davidson (1987). Dash–dot line encloses field of calc-alkaline rocks of the central Lesser Antilles, the continuous
line that of low-K tholeiites from the northern islands. Dashed lines marked FX indicate that fractional crystallization has little effect on Rb/K
ratios, unless amphibole is part of the fractionating assemblage [FX (am)]. The mixing line illustrates the effects of crustal contamination by a
high Rb/K component. Soufriere eruptive rocks occupy a transitional position between the fields of low-K tholeiites and calc-alkaline rocks (see
Fig. 17), but the relatively constant Rb/K ratios do not indicate a significant role for assimilation and fractional crystallization (AFC) in the
evolution of the Soufriere suite.
crystallization. The enrichment factors for the elements
most likely to be strongly incompatible, such as Ba, Cs
and Rb, are some 4–5 (Fig. 14b), which is consistent
with the 76 wt % crystallization figure found in the major
element modelling.
Although trace element distribution is generally consistent with evolution of the suite by closed-system fractional crystallization, the LREE depletion relative to
MREE and HREE during fractionation from basalts to
basaltic andesites creates a problem (Fig. 15). We have
modelled variations in La/Yb ratios by Rayleigh fractionation, using, as proxies for phenocryst partition coefficient, data for mineral phases (and glass) separated
from a Soufriere cumulate xenolith by Dostal et al. (1983).
The fractionating assemblages are those estimated from
major element modelling. The calculated trend shows
the expected LREE enrichment, from [La/Yb]N = 1·63
at 12·5 wt % MgO to 1·80 at 2·3 wt % MgO, and does
not provide a ready explanation for the observed LREE
depletion.
We noted earlier that the Soufriere rocks represent
several magmatic lineages. Rather than a crystal fractionation control, the LREE depletion in more evolved
rocks may reflect the fact that these rocks were derived
from parental magmas with lower LREE/HREE ratios
than the magnesian rocks in our analysed suite. This
might also explain the TiO2 decrease, followed by an
increase, in increasingly evolved rocks.
A role for amphibole?
Amphibole is a common phenocryst phase in several
Lesser Antilles suites, where it is restricted to evolved
rocks (andesites and dacites), e.g. the low-K tholeiitic
series of Mt Misery (now Mt Liamuiga) volcano, St Kitts
(Baker, 1968), the calc-alkaline rocks of Montserrat (Rea,
1974) and Mt Pelée, Martinique (Smith & Roobol, 1990)
and the M-series association of Grenada (Devine, 1995).
It has not, however, been found as a phenocryst phase
in any Soufriere St Vincent rock. The occurrence of
amphibole in cumulate blocks and as occasional xenocrysts in the 1979 eruption products may, however, be
evidence of high magmatic water contents at depth, and
the question is raised whether all Soufriere magmas went
through a higher-pressure phase of amphibole crystallization, which was followed by amphibole resorption
during ascent. Compositional variation in the eruptive
suite may therefore reflect a cryptic amphibole control,
1748
HEATH et al.
MAGMAGENESIS AT SOUFRIERE VOLCANO
as has been suggested for other calc-alkaline suites (Foden
& Green, 1992; Miller et al., 1992; Romnick et al., 1992).
Amphibole stability is dependent on temperature. Even
at high pressures and water contents, amphibole is unlikely to be stable above 1000°C (Helz, 1982; Allen &
Boettcher, 1983; Spulber & Rutherford, 1983), with the
possible exception of fluor-richterites, which might be
stable above 1200°C (Arculus, 1994). The basalts and
basaltic andesites of Soufriere St Vincent may simply
have been too hot (T [1050°C) to crystallize amphibole
as a near-liquidus phase.
Amphibole stability is also sensitive to melt composition. Cawthorn & O’Hara (1976) and Cawthorn
(1976) showed that in the system CMAS–Na2O–H2O at
pH2O = 5 kbar, hydrous basaltic melts needed [3 wt
% Na2O to crystallize pargasitic amphibole. Sisson &
Grove (1993a) confirmed, from melting experiments on
natural high-alumina basalts at 2 kbar (water saturated),
that hornblende stability is sensitive to Na2O content, as
well as H2O content, of the melt, and pointed out that
although hornblende can form as a (near-)liquidus phase
in hydrous, Na-rich magmas, it will not be stable in
moderate-Na systems until melt compositions are andesitic. Na2O concentrations [3 wt % tend to be restricted,
in the Lesser Antilles, to rocks with Ζ5 wt % MgO; the
restriction of amphibole to more evolved rocks is thus in
line with the experimental evidence.
In addition to the experimental evidence, we showed
earlier that production of basalts with MgO ~8 wt %
from more magnesian (MgO = 12·5 wt %) parental
basalts at Soufriere could be satisfactorily modelled using
combinations
of
the
crystallizing
assemblage
ol + cpx + pl + Cr-sp + Ti-mag + opx. Nevertheless,
an equally good fit (Rr 2 = 0·03), with about the same
inferred degree of crystallization, is achieved using the
assemblage 10% ol + 4% cpx + 11% pl + 0·1% Crsp + 17% hb, which is matched by assemblages in cumulate blocks (Arculus & Wills, 1980). However, adding
amphibole to the high-pressure crystallizing assemblage
has little effect on trace element distribution, as likely
partition coefficients, except for Ba and Rb, are similar to
those for clinopyroxene (Dostal et al., 1983). In particular,
amphibole does not provide a solution to the problem
of LREE depletion relative to HREE, as partition coefficients for HREE are much larger than those for LREE
(Sisson, 1994).
We suggest that amphibole was not involved in the
evolution of the Soufriere suite for the following reasons:
(1) there is a compositional and mineralogical continuum
from magnesian basalts, last equilibrated in the mantle,
to andesites; amphibole is not present as a phenocryst
phase; (2) major element modelling indicates that compositional variation in the suite can be explained without
recourse to amphibole fractionation; (3) the magmas may
simply have been too hot and too Na poor to stabilize
amphibole.
DISCUSSION
A summary of the evolutionary sequence of Soufriere
magmas is shown schematically in Fig. 25. By analogy
with magmatism on Grenada, the most primitive magmas
may have been derived from more olivine-rich (picritic)
magmas; however, their compositions are equally compatible with their being primary magmas in their own
right. These most primitive Soufriere magmas may have
been generated by ~15% partial melting of a mantle
source which resembled the N-MORB source before
addition of a subduction-related component which contained contributions from subducted sediments and basaltic crust. The most primitive magmas equilibrated
with mantle at 50–60 km depth; temperatures were
around 1130° and f (O2) was more oxidizing than
FMQ + 1. A period of olivine + spinel ± clinopyroxene
fractionation was followed by the appearance of plagioclase as a phenocryst phase. The pre-eruptive water
contents of these magmas are uncertain but the levels
were insufficient to prevent relatively early crystallization
of
plagioclase,
such
that
the
assemblage
ol + sp + cpx + plag is found in rocks with MgO >10
wt %. Fractionation of this assemblage resulted in modest
increases in Al2O3 content and FeO∗/MgO ratios of
residual liquids, that is, in a trend transitional between
tholeiitic and calc-alkaline.
In the early stages of Soufriere’s evolution, some of
the more primitive magmas reached the surface relatively
unmodified; rocks with the highest mg-numbers are restricted to the earliest stratigraphic unit. As the plumbing
system matured, magmas tended to reside for longer
periods at the base of, and within, the crust and the most
recent products have tended to be basaltic andesites (T
~1050°C). Judging from our pressure estimates (Figs 7
and 8), magma reservoirs existed at depths of ~30 km
(equivalent to 1 GPa) when ol + sp + cpx + plag were
the fractionating phases and magma compositions
evolved through towards basaltic andesite. The depth of
higher-level magma reservoirs is difficult to determine
without an accurate knowledge of magmatic water contents. However, basaltic andesites may have resided at
mid-crustal depths (10–15 km; Figs 7 and 8). This estimate
is consistent with ground tilt measurements which, over
a 12 year period including the 1979 eruption, were
interpreted to point to the presence of a magma chamber
at a depth of >10 km (Fiske & Shepherd, 1982, 1990).
Evidence relating to the water contents of the (basaltic)
andesite magmas is equivocal but is consistent with the
possibility that they were hydrous but water undersaturated.
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VOLUME 39
NUMBER 10
OCTOBER 1998
Fig. 25. Schematic model for the evolution of Soufriere magmas. The total number and dimensions of reservoirs are unknown, but those shown
here are consistent with the P–T data of this study.
There is no isotopic or trace element evidence for
significant crustal contamination of the Soufriere
magmas, despite the ubiquity of the process in other
calc-alkaline suites in the Lesser Antilles. The efficiency
of AFC processes is related to such factors as the rate of
magma supply, the temperature gradient between magma
and wall rocks, and the thickness of crystal mush on the
chamber walls and roof. Soufriere eruptive products carry
unusually abundant cumulate nodules (Arculus & Wills,
1980), which may indicate thick, marginal crystal deposits. These in turn may have armoured the magmas
from reaction with, and contamination by, the country
rocks. The development of thick marginal facies may be
a result of low magma supply rates at Soufriere, which
would be consistent with the relative rarity of mixed
magma rocks. Data for volumetric volcanic production
for the last 0·1 my show that the islands with the strongest
isotopic evidence for crustal contamination have had the
highest volumetric production (>5 km3; Wadge, 1986,
fig. 10).
In addition, volcanic activity may have started on St
Vincent much more recently than on neighbouring
islands; in contrast to St Lucia, Martinique, Grenada
and many of the Grenadine islands, St Vincent lacks
exposed Miocene rocks. The lack of deposits older than
~2·8 Ma may simply reflect incomplete sampling or
complete burial by post-Pliocene volcanic deposits. However, there is some support for the relative youth of St
Vincent from seismic data, as discussed by Wadge (1986).
The seismic refraction horizon between the upper- and
middle-crustal layers rises beneath most of the Lesser
Antillean islands, which has been interpreted to reflect
large volumes of intruded material. Wadge (1986) suggested that large andesitic plutons underlying the islands
usually intercept rising basaltic magmas, which then
evolve towards andesitic and dacitic compositions
through AFC and mixing. However, the seismic boundary is flatter and less distinct beneath St Vincent, which
may indicate a briefer history of magmatism, and less
well-developed magma reservoirs, which would be con-
1750
HEATH et al.
MAGMAGENESIS AT SOUFRIERE VOLCANO
sistent with the relative abundance of basaltic rocks and
limited amounts of mixing and contamination.
We suggest that the parental magmas to other calcalkaline suites in the Lesser Antilles were similar to the
magnesian basalts of Soufriere and that the development
of a more strongly calc-alkaline character in evolved
rocks was a result of AFC processes in the crust resulting
from higher magma supply rates and more mature magmatic plumbing systems.
ACKNOWLEDGEMENTS
We thank Richard Robertson of the Seismic Research
Unit, University of the West Indies, for invaluable help
in the field. Isotopic analyses were carried out by E.H.
at the Open University; the untiring assistance of Mabs
Johnston, Peter van Calsteren and Simon Turner is
gratefully acknowledged. Thanks are also due to Nick
Rogers (Open University) for INAA data, and Vicky
Burnett, Paul Wheeler and Stuart Black (Lancaster University) for help with XRF analyses. We thank the Natural
Environment Research Council for support, through a
research studentship to E.H. R.M. and H.S. thank NATO
for support for field-work on St Vincent. Finally, we
thank Drs S. Eggins and I. Parkinson for very helpful
journal reviews.
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JOURNAL OF PETROLOGY
VOLUME 39
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APPENDIX A: SAMPLE
DESCRIPTIONS
Abbreviations used: BA, basaltic andesite; PSL, PreSomma Lava; YTF, Yellow Tuff Formation; CL, Crater
Lava; PF, Pyroclastic Formation.
STV 301: 13°15.65′N, 61°07·05′W. Basaltic PSL from
Black Point.
STV 303: 13°15′N, 61°07′W. Basaltic pumice from YTF
fall deposit, near Mangrove.
STV 306: 13°18·01′N, 61°07·24′W. BA scoria from PF,
lower Rabacca.
STV 307: 13°18·07′N, 61°07·54′W. BA pumice from
lapilli fall layer in YTF, lower Rabacca.
STV 309: 13°18·3′N, 61°08·5′W. PSL block from flow
breccia overlain by YTF, Rabacca River, below Chainspout.
STV 310: 13°18·2′N, 61°08·3′W. Basaltic PSL from
Rabacca River, below Chainspout.
STV 312: 13°18·45′N, 61°07·14′W. BA scoria in PF
pyroclastic flow, Waterloo sea cliff, Oven Gully.
STV 313: 13°20·33′N, 61°07·40′W. BA pumice from
grey fall unit close to base of roadside YTF succession,
Robin Rock.
STV 314: 13°20·2′N, 61°07·4′W. BA PSL from flow
underlying YTF, Robin Rock Point.
STV 315: 13°20·59′N, 61°07·66′W. Basaltic PSL, London Point.
STV 316: 13°21.21′N, 61°07·87′W. BA PSL, Chibarabu
Point.
STV 317: 13°21·38′N, 61°07·97′W. Basaltic scoria from
YTF, Sion Hill.
STV 318: 13°22·52′N, 61°08·45′W. BA PSL from Cowand-calves flow, Owia salt pond.
STV 319: 13°22·59′N, 61°09·24′W. Basaltic PSL from
Commantawana Bay.
STV 320: 13°22·6′N, 61°09·3′W. Basaltic scoria from
>1 m thick YTF fall, Jumby Point.
STV 323: 13°21·54′N, 61°07·89′W. BA PSL, south
Sandy Bay.
STV 324: 13°19·66′N, 61°10·71′W. BA Crater Lava,
Soufriere Table flow.
STV 325: 13°19·82′N, 61°10·56′W. BA pumice from
1979 fall deposit, Windward Trail.
STV 326: 13°20·8′N, 61°10·75′W. BA PSL, basal flow
from Soufriere Mountains.
STV 330: 13°18·52′N, 61°07·87′W. BA pumice from
lowest fall layer, Lot 14.
STV 331: 13°18·52′N, 61°07·87′W. BA pumice from
(1718?) fall layer, Lot 14.
STV 332: 13°18·52′N, 61°07·87′W. BA pumice from
(1812?) eruption, Lot 14.
NUMBER 10
OCTOBER 1998
STV 333: 13°18·997′N, 61°09·169′W. Basaltic crystalrich fall unit, pre-1718, Bamboo.
STV 334: 13°18·64′N, 61°08·81′W. Basaltic scoria clasts
from 1902 surge deposit? Lot 14.
STV 335: 13°18·64′N, 61°08·81′W. BA tephra clasts
from fall deposit (1812?), Lot 14.
STV 345: 13°17·44′N, 61°08·27′W. BA PSL from
quarry, Indian Estate.
STV 349: 13°18·68′N, 61°07·99′W. BA scoria from PF
pyroclastic flow, Waribishy River.
STV 351: 13°18·25′N, 61°08·43′W. BA PSL, lowermost
lava near road, Waribishy.
STV 354: 13°18·721′N, 61°09·795′W. BA scoria from
PF, upper Rabacca River.
STV 356: 13°18·8′N, 61°09·9′W. Basaltic PSL, lower
flow at waterfall cliff, top of valley, upper Rabacca River.
STV 357: 13°18·8′N, 61°10·0′W. BA PSL, upper cliffforming flow, upper Rabacca River.
STV 358: 13°18·25′N, 61°08·43′W. Basaltic PSL, lava
bed at falls, Rabacca River.
STV 359: 13°17·969′N, 61°08·794′W. BA scoria, 1902
pyroclastic flow, Dry Rabacca.
STV 362: 13°19·05′N, 61°13·57′W. BA scoria from
1902? pyroclastic flow, south bank of Dry Wallibou.
STV 363: 13°19′N, 61°13′W. BA scoria from PF pyroclastic flow, Dry Wallibou.
STV 365: 13°19′N, 61°13′W. BA banded scoria from
PF pyroclastic flow, Wallibou–Ronde.
STV 369: 13°22·2′N, 61°11·7′W. BA PSL from upper
flow near concrete bridge, Baleine Falls.
STV 371: 13°20·3′N, 61°13·0′W. BA scoria from PF
pyroclastic flow, Larikai beach.
STV 372: 13°20·2′N, 61°12·7′W. BA PSL, platy lava
beneath YTF, Larikai River.
STV 373: 13°17·3′N, 61°14·4′W. BA pumice from fall
unit in YTF, south Chateaubelair.
STV 374: 13°17·2′N, 61°15·1′W. Basaltic lapilli from
fall layer in YTF, Dark View.
STV 376: 13°17·1′N, 61°15·2′W. Pumice clasts from
mixed YTF fall unit, Rose Bank. Separated into dark
basaltic (D) and light andesitic (L) components.
STV 377: 13°17·1′N, 61°15·4′W. Banded pumice clasts
from mixed YTF fall unit, Rose Bank. Separated into
dark BA (D), intermediate BA (I) and light andesitic (L)
components.
STV 79-70: 13°20′N, 61°11′W. BA lava from 1979
crater dome.
SVE 113: 13°20′N, 61°11′W. BA CL from thickest flow
in NW (Larikai) wall of crater.
SVE 114: 13°20′N, 61°11′W. Andesitic CL from Nose
Dike, cutting SE interior of crater wall.
SVE 115: 13°20′N, 61°11′W. BA CL, 15 m thick lava,
SW wall of crater.
SVE 116: 13°20′N, 61°11′W. BA CL, prominent in W,
S and SW of crater wall.
1754
HEATH et al.
MAGMAGENESIS AT SOUFRIERE VOLCANO
Table A1: Modal proportions (vol. %) of phenocrysts in selected Soufriere rocks, in order of increasing silica
content
Sample no.
Recalc’d
Vol. % of total phenocrysts (vesicle- and groundmass-free)
Vesic.-free
loss-free
Vesic. (%)
grdms (%)
SiO2 (wt %)
Plag %
Cpx %
STV 301
47·2
trace
19·1
STV 358
47·5
13·2
14·1
STV 310
47·7
26·6
STV 356
48·0
STV 334
Opx %
Oliv %
Ox %
0
67·7
13·2
0
67·5
0
0
67·7
5
0
53·9
0
22·6
0
50
0·8
0
60·6
0
14·7
17·4
0
61
6·9
0
66·7
0
50·9
63·0
15·8
0
21·2
trace
0
58·1
32·3
STV 315
51·7
69·6
7·8
trace
17·0
4·6
1·0
33·0
0
STV 319
52·3
70·2
23·0
1·1
3·9
1·9
0
43·0
0·5
STV 309
52·5
62·3
11·5
0·6
25·1
0·6
0
64·5
8·9
STV 324
52·6
70·1
7·1
1·0
19·3
2·5
0
39·5
2·4
STV 320
52·7
59·3
19·8
2·5
17·1
1·2
0
51·6
28·1
STV 303
53·0
67·8
15·3
7·9
1·4
7·6
0
51·6
18·1
STV 363
53·1
64·5
13·2
10·3
2·1
4·6
5·3
52·4
13·9
STV 373
53·4
72·2
16·7
4·6
1·5
5·0
0
53·9
19·1
STV 306
53·7
69·7
8·7
9·2
5·4
7·0
0
45·9
44·5
STV 330
54·0
70·5
9·1
5·9
5·9
8·6
0
52·5
29·7
STV 354
54·1
57·8
10·8
9·1
1·1
7·0
14·2
52·8
16·1
STV 79-70
54·2
71·0
9·1
9·8
1·7
8·4
0
42·7
23·5
STV 326
54·3
63·5
10·8
14·6
5·2
6·0
0
48·0
0
STV 362
54·4
67·8
12·3
3·6
3·8
6·1
6·3
50·6
42·9
STV 372
54·4
69·1
12·3
4·3
5·5
5·9
2·9
48·8
0
STV 312
54·8
66·1
19·7
9·2
trace
5·0
0
50·2
13·2
STV 307
55·0
76·3
8·1
5·2
3·9
6·5
0
61·6
25·8
STV 371
55·1
63·5
14·1
11·1
3·6
7·7
0
55·9
18·0
STV 323
55·2
69·8
13·5
1·6
4·2
8·7
2·2
45·0
7·6
STV 351
55·2
79·9
8·5
5·2
3·1
3·3
0
54·2
7·1
STV 318
55·4
71·0
15·3
1·8
trace
11·9
0
39·4
0
STV 345
55·5
62·1
14·5
9·3
8·9
5·2
0
53·8
0
STV 314
55·7
73·5
14·6
3·7
2·6
5·6
0
46·4
16·3
STV 369
55·8
73·7
14·4
2·0
2·1
7·8
0
43·8
0
SVE 113
56·0
59·2
24·3
5·4
trace
7·7
3·3
42·9
0
STV 376(L)
61·9
79·9
8·1
4·9
0
7·1
0
71·7
30·5
1755
Xen %
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 10
OCTOBER 1998
Table A2: Major element abundances (wt % oxides) for Soufriere rocks, determined by XRF spectrometry
Sample no.
SiO2
Al2O3
Fe2O3∗
CaO
MgO
MnO
Na2O
K2O
TiO2
P2O5
Balance
Pre-Somma Lavas
STV 301
47·01
15·28
9·77
10·96
12·50
0·16
2·23
0·47
1·07
0·12
0·43
STV 309
51·55
17·02
8·67
9·14
7·72
0·17
2·74
0·32
0·80
0·11
1·76
STV 310
47·29
15·60
9·83
10·48
12·21
0·17
2·22
0·26
0·94
0·10
0·91
STV 314
54·98
17·75
8·48
8·22
3·72
0·17
3·46
0·72
0·96
0·13
1·41
STV 315
51·20
16·88
8·95
10·55
7·76
0·16
2·33
0·33
0·82
0·09
0·92
STV 316
55·12
17·57
8·46
8·56
4·26
0·18
3·43
0·57
0·84
0·14
0·87
STV 318
55·14
17·66
8·97
8·20
3·63
0·19
3·91
0·68
0·97
0·14
0·50
STV 319
51·68
17·82
8·86
10·43
5·67
0·16
2·57
0·56
0·82
0·11
1·31
STV 323
54·89
17·64
8·87
8·57
4·47
0·18
3·28
0·54
0·83
0·13
0·60
STV 326
53·79
17·42
8·93
8·72
5·20
0·18
3·15
0·54
0·92
0·14
1·01
STV 345
55·03
17·46
7·95
8·42
5·42
0·15
3·23
0·48
0·80
0·14
0·92
STV 356
47·33
16·18
9·93
10·44
11·09
0·17
2·24
0·29
0·90
0·10
1·34
STV 351
54·76
19·59
7·87
8·69
2·71
0·16
3·62
0·65
0·97
0·15
0·83
STV 357
54·78
18·68
8·71
8·88
3·78
0·18
3·17
0·66
0·86
0·14
0·16
STV 358
47·29
15·71
10·00
10·32
12·53
0·17
2·25
0·28
0·94
0·10
0·41
STV 369
54·63
18·09
8·40
7·99
3·18
0·16
3·65
0·68
0·98
0·16
2·09
STV 372
53·24
17·65
8·38
8·76
5·05
0·17
2·98
0·54
0·87
0·13
2·23
Yellow Tuff Formation
STV 303
51·43
17·97
8·97
9·20
5·23
0·18
2·59
0·40
0·91
0·12
2·99
STV 307
54·31
18·20
8·44
8·65
4·22
0·18
3·15
0·56
0·91
0·13
1·26
STV 313
55·72
17·98
8·43
8·09
3·68
0·18
3·29
0·70
0·89
0·13
0·93
STV 317
51·92
18·99
8·64
9·89
4·62
0·17
2·76
0·51
0·88
0·11
1·51
STV 320
51·50
18·44
9·23
9·21
4·50
0·18
3·10
0·45
1·01
0·11
2·27
STV 373
52·88
18·52
9·05
9·39
4·51
0·18
2·88
0·48
0·93
0·12
1·06
STV 374
50·06
15·56
9·03
9·73
10·37
0·16
2·36
0·43
0·85
0·12
1·32
STV 376(L)
59·64
16·56
6·18
5·82
2·31
0·15
3·92
0·79
0·60
0·17
3·86
STV 376(D)
51·58
16·03
8·21
9·81
8·66
0·15
2·62
0·46
0·77
0·11
1·59
STV 377(L)
58·07
16·73
6·48
6·52
3·32
0·15
3·68
0·64
0·61
0·16
3·62
STV 377(I)
55·39
16·61
7·37
7·77
5·55
0·16
3·23
0·54
0·68
0·14
2·56
STV 377(D)
53·76
16·45
7·64
8·26
6·42
0·16
3·09
0·49
0·69
0·14
2·90
STV 324
52·18
19·23
8·25
10·04
5·08
0·16
2·86
0·46
0·81
0·11
0·81
SVE 113
54·83
17·55
8·33
8·16
3·89
0·17
3·37
0·56
0·90
0·13
2·11
SVE 114
56·69
17·56
8·42
6·49
3·23
0·18
3·68
0·68
0·96
0·14
1·97
SVE 115
56·16
17·93
8·91
6·82
3·29
0·21
3·50
0·59
0·89
0·15
1·55
SVE 116
55·59
17·75
8·73
8·18
3·93
0·19
3·38
0·54
0·91
0·14
0·66
Crater Lavas
1756
HEATH et al.
Sample no.
SiO2
Al2O3
Fe2O3∗
MAGMAGENESIS AT SOUFRIERE VOLCANO
CaO
MgO
MnO
Na2O
K2O
TiO2
P2O5
Balance
Pyroclastic Formation
STV 306
53·40
18·15
9·08
9·28
4·52
0·18
3·14
0·55
0·95
0·12
0·63
STV 312
54·08
18·09
8·84
8·71
3·94
0·18
3·25
0·58
0·90
0·13
1·29
STV 325
54·04
18·33
9·09
9·02
4·20
0·18
3·25
0·52
0·99
0·13
0·25
STV 330
53·61
18·53
9·29
8·95
4·32
0·19
2·88
0·46
0·96
0·11
0·69
STV 331
54·37
18·30
8·91
8·60
3·99
0·19
3·07
0·52
0·91
0·14
1·00
STV 332
54·15
18·10
9·00
8·55
3·62
0·19
3·23
0·51
0·92
0·13
1·60
STV 333
50·01
17·73
8·94
9·75
7·89
0·16
2·42
0·36
0·89
0·10
1·77
STV 334
50·54
17·03
9·49
10·24
7·95
0·17
2·50
0·40
0·92
0·11
0·66
STV 335
54·87
17·97
9·11
8·30
3·55
0·20
3·38
0·52
0·95
0·13
1·02
STV 349
53·53
18·36
8·87
9·01
4·19
0·18
3·23
0·54
0·96
0·12
1·00
STV 354
52·90
17·95
8·88
8·86
4·07
0·18
3·19
0·53
0·95
0·13
2·37
STV 359
53·98
18·43
9·13
8·92
4·22
0·19
3·14
0·49
0·96
0·12
0·42
STV 362
54·01
18·35
9·04
8·84
4·08
0·19
3·16
0·49
0·93
0·12
0·79
STV 363
52·14
18·90
9·32
9·19
4·05
0·19
2·96
0·40
0·90
0·13
1·81
STV 365
53·42
18·08
8·68
8·53
3·44
0·19
3·22
0·52
0·88
0·13
2·92
STV 371
54·81
18·47
8·69
8·78
3·70
0·19
3·25
0·50
0·88
0·13
0·61
STV 79-70
53·63
18·13
9·12
8·84
4·23
0·18
3·08
0·53
0·97
0·13
1·16
V71013
54·56
18·25
8·70
7·35
3·97
0·19
3·28
0·56
0·95
0·14
2·05
STV 336
52·12
10·21
5·11
25·09
2·24
0·16
1·18
1·92
0·40
0·13
1·43
STV 337
57·72
19·92
5·70
8·30
2·96
0·18
3·42
0·14
0·59
0·10
0·97
STV 339
54·68
16·97
6·00
7·08
4·11
0·18
2·86
0·38
1·03
0·13
6·57
STV 340
46·54
11·85
5·71
30·61
3·06
0·20
0·57
0·14
1·34
STV 353
50·16
14·94
9·09
10·67
11·41
0·17
0·72
0·07
1·03
Xenoliths
1757
1·54
0·20
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 10
OCTOBER 1998
Table A3: Trace element abundances (ppm) in Soufriere rocks, determined by XRF spectrometry (except Th
and U, determined by TIMS)
Sample no.
Ba
Cr
Nb
Ni
Rb
Sc
Sr
V
Zr
Y
Th
U
Pre-Somma Lavas
STV 301
74
728
4·7
250
7
41
202
299
61
18
0·849
0·348
STV 309
117
360
3·5
139
10
30
200
215
72
18
0·901
0·461
STV 310
75
672
3·9
289
6
37
192
269
54
17
STV 314
148
18
0·4
20
16
29
207
217
86
19
STV 315
81
303
2·9
75
8
39
167
277
51
18
0·493
0·263
STV 316
134
61
3·7
32
12
28
202
207
77
19
STV 318
157
22
3·9
19
15
30
206
220
80
19
STV 319
102
64
3·0
37
9
38
341
270
71
18
0·909
0·369
STV 323
133
69
3·5
32
13
27
201
200
81
19
1·063
0·515
STV 326
131
116
3·8
53
13
31
209
231
79
19
STV 345
133
223
4·0
86
13
25
238
181
91
19
1·025
0·510
STV 351
141
19
4·0
16
16
27
249
206
95
20
1·311
0·623
STV 356
56
663
2·9
213
7
39
205
281
58
18
STV 357
143
43
3·7
21
15
30
216
224
83
19
STV 358
66
654
4·0
301
5
34
194
248
59
18
0·562
0·284
STV 369
159
25
4·0
20
15
34
227
255
94
20
STV 372
135
119
3·8
56
14
29
216
218
80
19
1·225
0·590
0·959
0·557
Yellow Tuff Formation
STV 303
128
116
3·5
45
12
31
214
227
67
18
STV 307
136
34
3·7
21
14
32
208
233
77
19
STV 313
182
21
4·1
19
18
30
213
219
87
19
STV 317
109
42
3·1
34
11
33
247
243
67
18
STV 320
102
33
3·4
22
11
34
197
256
71
19
STV 373
125
34
3·0
22
13
31
215
230
71
19
0·981
0·475
STV 374
112
719
4·4
234
12
34
197
251
71
18
1·360
0·698
STV 376(L)
224
24
5·5
26
28
14
237
93
131
21
1·868
1·138
STV 376(D)
123
557
3·7
178
14
35
211
251
71
18
STV 377(L)
186
102
4·6
51
21
17
258
122
119
20
STV 377(I)
159
233
4·7
107
18
21
240
151
101
19
STV 377(D)
143
331
4·6
134
15
24
234
173
93
19
STV 324
106
126
3·1
49
12
28
220
204
72
19
0·957
0·461
SVE 113
133
64
4·0
26
14
29
214
212
83
19
SVE 114
22
3·5
18
SVE 115
22
3·6
18
52
4·0
195
83
19
Crater Lavas
SVE 116
124
22
12
190
193
26
1758
212
HEATH et al.
Sample no.
Ba
Cr
Nb
MAGMAGENESIS AT SOUFRIERE VOLCANO
Ni
Rb
Sc
Sr
V
Zr
Y
Th
U
1·271
0·599
0·980
0·486
1·077
0·534
Pyroclastic Formation
STV 306
129
48
4·0
16
14
31
213
230
73
19
STV 312
140
18
3·8
17
15
27
214
200
77
19
STV 325
122
35
3·7
24
13
30
206
223
77
19
STV 330
125
25
3·4
17
12
32
209
235
78
19
STV 331
135
20
3·4
20
14
28
226
210
84
19
STV 332
133
22
3·6
15
13
29
219
219
81
19
STV 333
103
302
3·0
116
11
39
195
283
63
18
STV 334
98
300
3·7
109
11
40
206
295
66
18
STV 335
128
10
3·7
14
14
30
208
221
81
19
STV 349
136
28
3·8
19
15
32
231
241
82
19
STV 354
133
29
3·8
22
14
33
212
248
81
19
STV 359
115
38
3·7
21
13
29
207
215
75
19
STV 362
126
31
3·8
21
14
31
218
232
80
19
STV 363
93
22
3·4
20
11
31
222
232
74
19
STV 365
136
13
4·0
18
14
30
233
227
87
19
STV 371
110
8
3·4
14
13
29
216
216
78
19
0·871
0·443
STV 79-70
117
40
4·1
23
13
31
208
225
77
19
1·077
0·532
43
3·4
V71013
17
197
Xenoliths
STV 336
334
42
4·4
31
43
25
386
151
81
19
STV 337
101
29
3·0
19
5
22
288
150
82
19
STV 339
120
27
3·8
20
15
36
220
275
92
21
STV 340
27
33
4·2
32
3
31
96
190
56
17
STV 353
139
681
2·4
252
12
44
292
308
61
18
1759
JOURNAL OF PETROLOGY
VOLUME 39
NUMBER 10
OCTOBER 1998
Table A4: Trace element abundances (ppm) for selected Soufriere rocks, determined by INAA
Sample no.
La
Ce
Nd
Sm
Eu
Tb
Yb
Lu
Ta
Hf
Cs
Pre-Somma Lavas
STV 301
5·7
12·7
8·9
2·47
0·92
0·56
1·74
0·27
STV 309
4·8
11·7
8·6
2·30
0·87
0·59
1·99
0·33
0·17
2·38
1·68
STV 315
3·2
8·2
8·3
2·24
0·83
0·58
2·37
0·39
0·10
1·75
STV 319
4·9
11·7
9·6
2·47
0·89
0·59
2·25
0·36
0·10
1·91
0·36
STV 323
5·7
14·2
11·4
3·07
1·05
0·74
2·98
0·46
0·18
2·35
0·49
STV 345
6·8
15·0
11·6
3·18
1·08
0·68
2·63
0·41
0·19
2·65
0·39
STV 351
6·2
15·6
11·3
3·27
1·20
0·76
3·13
0·48
0·23
2·69
0·64
STV 358
4·0
10·5
8·4
2·34
0·91
0·54
1·65
0·27
0·16
1·70
0·49
Yellow Tuff Formation
STV 303
5·1
12·1
8·6
2·68
0·95
0·68
2·47
0·39
0·13
2·13
STV 373
5·2
12·5
8·9
2·88
1·03
0·69
2·73
0·44
0·15
2·10
0·54
0·47
STV 374
5·7
13·4
10·0
2·60
0·94
0·61
2·20
0·34
0·21
2·09
0·59
STV 376(L)
8·7
20·2
13·4
3·16
0·99
0·64
2·71
0·44
0·30
3·60
1·18
5·2
12·7
10·0
2·74
0·99
0·69
2·55
0·41
0·15
2·10
0·29
0·69
Crater Lavas
STV 324
Pyroclastic Formation
STV 312
6·3
15·6
12·3
3·27
1·13
0·83
3·10
0·48
0·19
2·55
STV 334
4·9
12·1
8·8
2·54
0·95
0·62
2·21
0·36
0·16
1·88
0·39
STV 354
5·9
14·1
10·6
3·14
1·11
0·74
2·95
0·45
0·18
2·43
0·60
STV 371
5·1
12·9
10·0
2·98
1·08
0·71
2·98
0·47
0·13
2·37
0·51
STV 79-70
5·9
13·8
10·5
3·18
1·12
0·76
3·01
0·46
0·19
2·34
0·50
1760
HEATH et al.
MAGMAGENESIS AT SOUFRIERE VOLCANO
Table A5: Sr, Nd and Pb isotopic data for selected Soufriere rocks, determined by mass spectrometry
Sample no.
87
Sr/86Sr
±
143
Nd/144Nd ±
206
Pb/204Pb
±
207
Pb/204Pb
±
208
±
Pb/204Pb
Pre-Somma Lavas
STV 301
0·703822
10
0·512946
6
STV 309
0·704204
10
0·512918
7
19·310
0·003
15·720
0·002
38·814
0·006
STV 315
0·70391
10
0·512976
7
19·155
0·002
15·695
0·001
38·766
0·004
STV 319
0·703777
10
0·512843
5
18·980
<0·001
15·639
<0·001
38·631
0·001
STV 323
0·70411
10
19·143
0·001
15·657
0·001
38·657
0·003
STV 345
0·703941
10
0·512940
11
19·279
<0·001
15·709
<0·001
38·834
0·001
STV 351
0·704061
10
0·512883
4
STV 358
0·703752
10
19·160
0·001
15·689
0·001
38·718
0·002
STV 372
0·704061
10
19·258
0·002
15·716
0·002
38·797
0·004
19·328
<0·001
15·703
<0·001
38·818
0·001
19·227
<0·001
15·698
<0·001
38·823
0·001
Yellow Tuff Formation
STV 303
0·704303
10
0·512892
9
STV 373
0·704229
10
0·512864
6
STV 374
0·704407
10
0·512907
5
STV 376(L)
0·704254
10
0·703988
10
Crater Lavas
STV 324
0·512911
4
Pyroclastic Formation
STV 312
0·704253
10
19·171
<0·001
15·685
<0·001
38·764
0·001
STV 334
0·704338
10
19·326
0·001
15·722
0·001
38·908
0·001
STV 354
0·704209
11
0·512877
14
18·644
<0·001
15·625
<0·001
38·253
0·001
STV 371
0·704024
11
0·512926
10
19·077
<0·001
15·676
<0·001
38·683
0·001
STV 79-70
0·704118
10
0·512895
6
19·252
0·001
15·692
<0·001
38·813
0·001
STV 336
0·707745
10
STV 337
0·703843
10
STV 339
0·704061
10
STV 340
0·707172
10
0·512323
10
STV 353
0·706163
10
0·512895
8
Xenoliths
1761
2·16
2·57
2·98
2·88
2·70
2·47
2·88
323 c3 melt: midway in opx
323 d5 melt: midway in opx
323 e7 melt: midway in opx
323 e9 melt: midway in opx
323 f10 melt: core of opx
323 h12 melt: core of cpx
2·15
2·05
3·14
1·97
2·15
1·49
1·53
1·35
2·35
1·55
2·56
2·56
2·75
3·21
3·17
3·22
4·02
3·41
3·77
2·84
2·63
3·15
3·35
3·40
303 a1 melt: core of opx
303 a3 melt: core of opx
303 a4 melt: rim of opx
303 c6
303 c7
303 d8
303 d9
303 e10 melt: core of cpx
303 e11 melt: core of cpx
303 f13 melt: core of cpx
303 h15 melt: in titano-magnetite
303 h16 melt: in titano-magnetite
303 h17 melt: in titano-magnetite
376(L) a1 melt: midway in opx
376(L) a2 melt: midway in opx
376(L) a3 melt: midway in opx
376(L) b4 melt: core of opx
376(L) c5 melt: midway in plag
376(L) c6 melt: midway in plag
376(L) d7 melt: core of opx (in plag)
376(L) d8 melt: core of opx (in plag)
376(L) e9melt: core of opx
376(L) f11
376(L) f12
Yellow Tuff Formation
2·61
315 h12 melt: core of zoned cpx
Na2O
315 a2 melt: midway in cpx
Pre-Somma Lavas
Melt inclusion descriptions
1762
0·10
0·04
0·08
0·03
0·11
0·06
0·05
0·00
0·08
0·00
0·00
0·27
0·26
0·22
0·21
0·21
0·22
0·30
0·26
0·25
0·25
0·25
0·32
0·31
0·31
0·32
0·27
0·25
1·40
1·42
1·27
1·07
0·94
0·84
1·29
1·42
1·44
1·39
1·42
1·20
1·12
1·10
0·70
0·63
0·50
0·76
0·68
0·65
0·61
0·76
0·80
0·83
2·04
0·93
1·14
1·26
1·24
1·31
1·04
2·59
K2O
0·01
0·01
0·01
0·00
0·01
0·00
0·00
0·01
0·01
0·01
0·00
0·02
0·01
0·01
0·02
0·03
0·02
0·03
0·02
0·01
0·02
0·02
0·02
0·02
0·01
0·00
0·00
0·00
0·01
0·01
0·00
0·00
SO3
69·40
69·47
69·95
69·32
67·57
61·26
66·55
70·42
70·85
69·48
69·76
63·13
62·35
62·13
62·08
63·76
60·22
61·90
60·69
61·86
61·89
58·15
60·60
61·18
71·41
73·27
74·79
73·21
73·38
72·79
74·59
66·74
SiO2
0·07
0·10
0·10
0·11
0·11
0·04
0·13
0·14
0·13
0·12
0·10
0·14
0·14
0·17
0·12
0·12
0·15
0·21
0·27
0·12
0·15
0·30
0·19
0·16
0·06
0·01
0·02
0·09
0·09
0·05
0·07
0·07
MnO
0·10
0·08
0·03
0·06
0·08
0·09
0·16
0·09
0·10
0·05
0·09
0·47
0·52
0·48
0·31
0·23
0·25
0·20
0·19
0·29
0·29
0·26
0·27
0·28
0·34
0·25
0·29
0·24
0·25
0·33
0·27
0·34
P2O5
2·34
2·28
2·38
2·59
3·20
6·33
2·71
2·46
2·18
2·37
2·22
5·21
5·04
5·20
3·12
3·85
4·03
5·13
5·86
4·58
4·42
3·88
4·66
4·58
1·37
3·32
2·65
3·05
2·95
2·81
2·04
2·94
CaO
0·40
0·40
0·28
0·31
0·20
0·50
0·61
0·44
0·39
0·26
0·32
0·92
0·98
1·03
0·87
1·05
0·90
0·97
0·98
0·90
0·93
0·76
1·01
1·00
0·76
0·29
0·24
0·49
0·49
0·49
0·19
0·26
TiO2
2·47
2·49
3·12
2·95
3·14
2·70
3·81
3·05
2·98
3·44
3·20
5·74
5·92
5·96
3·61
3·99
4·98
4·87
5·52
6·04
5·73
8·27
6·32
6·28
2·60
1·65
1·85
2·14
2·17
1·98
1·33
1·45
FeO
0·46
0·48
0·42
0·40
0·45
0·65
1·08
0·46
0·49
0·44
0·44
1·75
1·82
1·85
0·69
0·55
0·90
0·98
1·25
1·26
1·25
5·40
1·65
1·57
0·02
0·25
0·11
0·10
0·06
0·07
0·04
1·00
MgO
0·02
0·00
0·05
0·03
0·03
0·04
0·00
0·04
0·01
0·02
0·06
0·04
0·01
0·05
0·04
0·00
0·00
0·00
0·05
0·04
0·00
0·05
0·00
0·02
0·07
0·03
0·00
0·00
0·05
0·02
0·02
0·06
BaO
13·98
13·93
13·77
13·89
14·95
19·60
14·26
14·26
13·54
13·54
13·66
17·65
17·38
17·34
18·23
18·21
17·12
16·87
16·59
17·03
16·93
13·40
16·63
16·81
18·24
16·97
15·72
17·50
17·13
17·02
16·04
15·22
Al2O3
0·10
0·07
0·08
0·06
0·09
0·08
0·09
0·06
0·09
0·06
0·06
0·08
0·07
0·09
0·07
0·06
0·06
0·06
0·07
0·06
0·06
0·04
0·07
0·05
0·11
0·07
0·06
0·07
0·09
0·10
0·06
0·15
94·30
94·23
94·75
93·54
93·74
96·03
94·26
97·00
95·56
94·46
94·65
99·25
98·08
98·14
91·60
94·97
90·62
93·65
93·81
95·11
94·40
94·51
94·40
95·05
100·16
99·63
99·69
101.16
101·03
99·71
97·99
93·76
–O=F, Cl Sum
NUMBER 10
0·01
0·00
0·04
0·00
0·00
0·00
0·28
0·29
0·25
0·25
0·17
0·23
0·23
0·43
0·22
0·25
0·25
0·25
0·25
0·26
0·61
Cl
VOLUME 39
0·00
0·00
0·00
0·00
0·00
0·04
0·00
0·03
0·05
0·01
0·02
0·09
0·10
0·00
0·03
F
Table A6: Compositional data (wt %) for melt inclusions in phenocrysts from Soufriere rocks, determined by electron
microprobe
JOURNAL OF PETROLOGY
OCTOBER 1998
2·53
3·22
2·78
3·19
3·26
2·56
2·28
376(L) g13 melt: midway in plag
376(L) g14 melt: midway in plag
376(L) h15 melt: midway in plag
376(L) h16 melt: midway in plag
376(L) h17 melt: midway in plag
376(L) i19 melt: core of cpx
376(L) i20 melt: midway in cpx
2·21
2·31
2·51
324 a2 melt: core of cpx
324 a3 melt: midway in cpx
324 e7 melt: midway in plag
1763
6·67
2·38
2·39
2·30
2·11
5·08
2·83
2·35
2·76
2·41
2·21
2·30
2·51
2·91
2·02
2·08
334A a2 melt: midway in cpx
334A a3 melt: midway in cpx
334A b5 melt: rim in ol
334A c6 melt: core of plag
334A d7 melt: core of plag
334A e9 melt: core of plag
334A e10 melt: core of plag
334A e11melt: core of plag
334A f12 melt: core of cpx
334A f13 melt: midway in cpx
334A f14 melt: midway in cpx
334A f15 melt: midway in cpx
371 a2 melt: midway in opx
371 g14 melt: midway in cpx
371 g15 melt: midway in cpx
6·14
312A g9 melt: in cpx
4·53
3·97
312A d4 melt: core of opx
312A g11 melt: in cpx
5·41
312A c3 melt: core of cpx
312A g10 melt: in cpx
8·34
312A a1 melt: core of cpx
Pyroclastic Formation
1·76
324 a1 melt: core of cpx
Crater Lavas
Na2O
Melt inclusion descriptions
0·05
0·03
0·00
0·05
0·00
0·02
0·00
0·04
0·00
0·00
0·02
0·01
0·00
0·01
0·06
0·16
0·03
0·00
0·05
0·13
0·00
0·00
0·04
0·03
0·00
0·00
0·02
0·00
0·00
0·00
0·00
0·08
F
0·30
0·28
0·23
0·19
0·22
0·19
0·17
0·13
0·16
0·12
0·07
0·21
0·14
0·23
0·15
0·26
0·18
0·13
0·20
0·26
0·07
0·20
0·26
0·25
0·41
0·27
0·27
0·23
0·21
0·26
0·22
0·19
Cl
0·65
0·76
0·84
0·80
0·84
0·71
0·76
1·06
1·01
0·86
0·43
1·15
0·75
0·75
0·73
3·40
0·95
0·74
3·07
2·75
0·16
4·41
1·35
1·40
1·59
1·16
1·16
1·44
1·41
1·20
1·40
1·26
K2O
0·02
0·01
0·03
0·05
0·02
0·02
0·03
0·00
0·03
0·02
0·01
0·01
0·03
0·02
0·03
0·03
0·08
0·01
0·02
0·03
0·01
0·00
0·00
0·00
0·01
0·00
0·01
0·00
0·00
0·01
0·00
0·00
SO3
68·91
68·97
70·71
61·07
61·17
58·17
58·26
57·46
55·13
56·34
56·37
59·92
55·85
58·57
56·11
68·32
64·63
63·68
65·53
67·01
69·15
75·52
75·51
76·37
76·94
65·82
67·80
69·19
69·69
66·43
69·93
68·50
SiO2
0·09
0·10
0·05
0·13
0·17
0·23
0·18
0·22
0·27
0·25
0·14
0·29
0·25
0·25
0·23
0·08
0·10
0·20
0·08
0·11
0·00
0·03
0·05
0·06
0·07
0·15
0·12
0·07
0·05
0·10
0·09
0·11
MnO
0·28
0·28
0·21
0·19
0·24
0·27
0·21
0·19
0·21
0·24
0·21
0·22
0·19
0·28
0·19
0·32
0·25
0·21
0·29
0·28
0·27
0·11
0·23
0·20
0·30
0·08
0·10
0·08
0·04
0·16
0·07
0·06
P2O5
3·15
2·96
4·45
4·66
4·95
5·99
5·37
6·60
7·10
6·91
8·47
5·05
9·07
5·77
6·61
3·40
3·85
3·81
4·16
3·31
2·19
0·66
1·90
1·66
1·01
3·19
2·58
2·40
2·29
2·59
2·33
2·14
CaO
0·95
0·97
0·71
1·40
1·02
1·85
1·72
1·80
1·78
1·87
1·37
1·82
1·57
1·39
1·65
0·72
0·93
0·85
1·03
0·92
0·57
0·93
0·43
0·44
0·96
0·40
0·33
0·32
0·24
0·54
0·35
0·31
TiO2
1·21
1·49
1·62
6·31
8·11
10·01
8·08
10·26
11·58
11·09
6·58
11·65
9·34
9·62
9·74
1·22
3·26
6·55
3·03
0·67
0·49
1·29
1·45
1·23
1·44
3·35
3·03
2·59
2·52
2·90
2·35
2·13
FeO
0·24
0·64
0·34
1·38
1·52
2·19
2·52
4·25
4·96
4·50
2·72
5·03
3·05
2·79
3·24
0·33
0·57
1·01
1·08
0·26
0·31
0·12
0·18
0·14
0·05
0·87
0·49
0·54
0·53
0·65
0·49
0·45
MgO
0·01
0·00
0·02
0·00
0·05
0·01
0·00
0·01
0·00
0·04
0·00
0·02
0·00
0·03
0·05
0·01
0·00
0·02
0·02
0·03
0·00
0·01
0·01
0·04
0·04
0·03
0·03
0·00
0·00
0·04
0·01
0·00
BaO
17·49
16·68
17·83
18·47
16·50
15·65
17·48
14·59
13·92
13·89
20·29
9·76
16·69
15·39
16·42
17·95
17·31
16·89
17·01
18·20
18·67
10·64
15·30
14·65
13·45
13·31
13·83
13·96
13·63
13·83
13·89
13·32
Al2O3
0·09
0·08
0·05
0·07
0·05
0·05
0·04
0·05
0·04
0·03
0·02
0·05
0·03
0·05
0·06
0·12
0·06
0·03
0·07
0·11
0·02
0·04
0·08
0·07
0·09
0·06
0·07
0·05
0·05
0·06
0·05
0·08
95·33
95·12
99·90
97·16
97·05
97·46
97·16
99·33
98·45
98·92
101·73
97·18
99·19
97·42
97·53
100·59
98·75
100·20
99·48
99·24
100·20
96·38
98·93
98·62
97·93
90·85
92·25
94·02
93·77
91·43
94·31
91·00
–O=F, Cl Sum
HEATH et al.
MAGMAGENESIS AT SOUFRIERE VOLCANO
JOURNAL OF PETROLOGY
VOLUME 39
APPENDIX B: ANALYTICAL
METHODS
Major and trace element compositions were measured
at Lancaster University using a Phillips 1400 XRF spectrometer, and are quoted in wt % and ppm, respectively.
Standards were run at frequent intervals and calibrations
were regularly drift corrected using a monitor buffer.
Relative 1r precision is <0·5% for most major elements
(~2% for Na2O, MnO and P2O5) and <3% for most
minor elements (~4% for Ba and Sc, and ~5·5% for Cr
and Nb). Concentrations of the REE plus Ta, Hf and
Cs were determined by INAA, and are quoted in ppm.
Samples were irradiated with neutron flux monitors and
standards at the Imperial College Reactor Centre, and
counted twice (to ensure good precision for both shortlived and longer-lived isotopes) by gamma-ray spectrometry at the Open University. Relative 1r precision
is <3·5% for all elements except Cs (8%) and Ta (6%).
Compositions of phenocrysts and silicate melt inclusions
were measured by electron microprobe at the US Geo-
NUMBER 10
OCTOBER 1998
logical Survey in Virginia. Sodium was analysed first to
minimize its loss. Major and minor elements were counted
for 20 and 40 or 60 s, respectively. Relative 1r precision
is estimated to be 1–2% for major elements and 5–10%
for minor elements. Sr, Nd and Pb isotopic ratios and
Th and U concentrations (in ppm) were determined
by solid-source thermal ionization mass spectrometry
(TIMS) using Finnegan MAT 261 and 262 instruments
at the Open University. Chemical separation procedures
for Sr/Nd, Pb and Th/U were undertaken in separate
clean chemistry laboratories, and total procedure blanks
indicated negligible contamination. Machine standards
and rock standards were also analysed to ensure the high
quality and reproducibility of the data. 87Sr/86Sr ratios
are relative to an average value of 0·710279 ± 19 for
the NBS 987 standard, and replicate analyses of the J&M
Nd standard yielded 143Nd/144Nd = 0·511794 ± 9. Pb
data were corrected for mass fractionation relative to the
NBS 981 standard. Average 1r uncertainties on Th
and U concentrations are ±0·005 and ±0·001 ppm,
respectively.
1764