JOURNAL OF PETROLOGY
VOLUME 52
NUMBER 11
PAGES 2187^2220
2011
doi:10.1093/petrology/egr044
An Investigation into the Nature of the Magmatic
Plumbing System at ParicutinVolcano, Mexico
MICHAEL C. ROWE1,2*, DAVID W. PEATE1 AND
INGRID UKSTINS PEATE1
1
DEPARTMENT OF GEOSCIENCE, UNIVERSITY OF IOWA, 121 TROWBRIDGE HALL, IOWA CITY, IA 52242, USA
2
SCHOOL OF EARTH AND ENVIRONMENTAL SCIENCES, WASHINGTON STATE UNIVERSITY, PULLMAN, WA 99164, USA
RECEIVED JUNE 23, 2010; ACCEPTED SEPTEMBER 1, 2011
The temporal evolution of erupted magma compositions at Paricutin
Volcano (Mexico) is often cited as a classic example of assimilation^fractional crystallization processes with significant progressive
changes in major element, trace element, and isotopic compositions
occurring over the relatively short 9 year lifespan of the volcano. In
this study, major and trace element compositions of olivine- and
orthopyroxene-hosted melt inclusions are integrated with new trace
element analyses of the erupted lavas and data for entrained xenoliths
and xenolith glasses to provide a more comprehensive evaluation of
the evolution of Paricutin Volcano that questions this view. Melt inclusion compositions are bimodal with an undegassed, low-Si population (Type I) similar in composition to the whole-rock samples
and a degassed, high-Si population (Type II) recording late-stage
degassing and crystallization of the magma. Despite the rapid
changes in lava composition, melt inclusions hosted in both olivine
and orthopyroxene do not record any progressive contamination or
mixing of magmas. Homogeneity of Type I melt inclusions within
single lava samples implies significant contamination prior to crystallization and potentially a decoupling of assimilation^fractional
crystallization processes. Pre-existing models of magma evolution at
Paricutin Volcano are not consistent with the melt inclusion results
or new trace element whole-rock data.Whole-rock and melt inclusion
trace element analyses corroborate previous studies, which have suggested that the early erupted material (Phase 1; February^July
1943) was of a compositionally distinct magma compared with the
bulk of the erupted material during Phase 2 (July 1943^1946).
There is a second compositional transition between the Phase 2 and
Phase 3 (1947^1952) lavas, marked by a sudden change in Zr/Nb
despite similar MgO values, that is consistent with the arrival of a
new magma batch. This transition occurs prior to the major
Paricutin volcano in Mexico is a relatively short-lived
(1943^1952) volcanic center (cinder cone þ associated lava
and tephra) that is often cited as the ‘classic’ example of
an assimilation and fractional crystallization (AFC) process. The composition of the Paricutin lavas and tephras
evolved over the course of the eruption from basaltic andesite to andesite (55^60 wt % SiO2: Fig. 1a; e.g. Wilcox,
1954), accompanied by an increase in 87Sr/86Sr (Fig. 1b)
and d18O that have been attributed to progressive crustal
assimilation of a single magma batch (McBirney et al.,
1987). Published models to explain the compositional variability in Paricutin lavas have required significant
amounts of crustal assimilation accompanying crystallization of olivine and plagioclase (Wilcox, 1954; McBirney
et al., 1987; Cebria¤ et al., 2011).
*Corresponding author. Telephone: (509) 335-6770.
E-mail: [email protected]
ß The Author 2011. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
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compositional change from basaltic andesite to andesite magmas in
the waning stages of the eruption that is consistent with progressive
crustal assimilation within this latest magma batch. These data
demonstrate that the petrogenetic evolution of magmas at Paricutin
is more complex than simple progressive assimilation and fractional
crystallization and requires the presence of three compositionally distinct magma batches at shallow levels.
trace element; magma chamber; melt inclusion; crustal
contamination; crystallization
KEY WORDS:
I N T RO D U C T I O N
JOURNAL OF PETROLOGY
VOLUME 52
NUMBER 11
NOVEMBER 2011
Fig. 1. (a) Eruption date vs SiO2 (whole-rock lava and tephra data); (b) inset plot of SiO2 vs 87Sr/86Sr. Data sources: Wilcox (1954); McBirney
et al. (1987); Luhr (2001); Cebria¤ et al. (2011).
Assimilation and crystallization are often inferred to be
intimately linked, with the latent heat of crystallization
providing the thermal driving force for continued crustal
assimilation (e.g. Bowen, 1928; DePaolo, 1981; Davidson &
Wilson, 1989; Kuritani et al., 2005). Most investigations
have been based largely on whole-rock geochemistry; however, analysis of melt inclusions trapped in the crystallizing
mineral phases could potentially provide a direct means
to test models of coupled assimilation^crystallization. A
comparison of compositional variations in primary melt
inclusions (trapped during crystallization) with their host
mineral composition should allow the compositional evolution of the magma during crystallization to be monitored, rather than relying on final homogenized bulk-rock
compositions. If assimilation and crystallization are a
coupled process, temporally and spatially related, it is expected that as crystallization proceeds and the host mineral composition evolves the melt inclusion compositions
will likewise change and record increased crustal contamination, as demonstrated by Kent et al. (2002) for Yemen
flood basalts. Numerous studies have evaluated thermal
and physical models for coupled assimilation and fractional crystallization in magma chambers, relating these
pre-eruptive physical processes to compositional variations
in the magma (e.g. McBirney et al., 1985; Spera &
Bohrson, 2001; Kaneko & Koyaguchi, 2004; Leitch, 2004;
Spera & Bohrson, 2004; Kuritani et al., 2005, 2007). Other
studies have suggested that assimilation can occur as a
rapid, late-stage process, potentially during the course of
an eruption (e.g. Dungan, 2005; Erlund et al., 2010) or
even that bulk lava compositions can record a different
petrogenetic history from melt inclusions, with lavas
recording deeper crystallization whereas the melt inclusions record shallow degassing-induced crystallization and
assimilation (e.g. Johnson et al., 2008).
The objective of this study is to test models of magma
evolution at Paricutin Volcano so as to develop a better
understanding of the progressive development of the magmatic plumbing system and the relative timing of crustal
assimilation and crystallization in evolving magma systems. We present compositional data for host minerals and
melt inclusions from a well-characterized suite of lavas
2188
ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
from throughout the eruptive history of Paricutin Volcano
(Wilcox, 1954; McBirney et al., 1987). In addition, we integrate new trace element analyses of the lavas and entrained crustal xenoliths and xenolith glasses with the
melt inclusion and crystal chemistry and literature data to
provide a more comprehensive view of the evolution of
the Paricutin magmatic system.
E RU P T I V E H I S T O RY O F
PA R I C U T I N VO L C A N O
Paricutin lies within the Michoaca¤n^Guanajuato volcanic
field (MGVF), which contains over 1000 small eruptive
centers (Appendix Fig. A1) over an 40 000 km2 region
(Hasenaka, 1994); there is no evidence for any previous volcanism at the site of the volcano. The 9 year eruption of
Paricutin began on 20 February 1943, following several
weeks of intensifying seismicity, and the total erupted
volume of basaltic andesite and andesite magma is estimated at 1·38 km3.
Previous studies have divided the eruption into phases
based on changes in eruptive behavior and magma composition (McBirney et al., 1987; Luhr, 2001; Pioli et al.,
2008). In the present study, we use a division into four
eruptive phases (1, 2, 3a, 3b), based predominantly on lava
compositions; these are a minor modification of the
McBirney et al. (1987) and Pioli et al. (2008) subdivisions.
Phase 1 includes material erupted from early to mid-1943;
this short period at the initiation of the eruption is
marked by a more primitive whole-rock composition and
lower K2O content in both lavas and tephra erupted prior
to July 1943 (Luhr, 2001; Pioli et al., 2008). Phase 2 extends
from July 1943 to 1946, a period during which the major
element and isotopic variations of the lavas were relatively
restricted (55 wt % SiO2; 87Sr/86Sr 0·7038; McBirney
et al., 1987). Together, Phases 1 and 2 make up most of the
erupted volume at Paricutin (75 vol. %; McBirney et al.,
1987). Crustal xenoliths were predominantly recovered
during the first 3 years of the eruption (during Phases 1
and 2), and comprise a variety of felsic igneous crustal
lithologies, variably altered and partially melted, consisting of feldspar and quartz crystals with varying proportions (up to 90 vol. %) of vesicular glass (McBirney et al.,
1987). Phase 3a (‘middle stage’ of McBirney et al., 1987) incorporates a rapid compositional shift from basaltic andesite to andesite during 1947 and early 1948 that accounts
for only 8^10% of the erupted volume. The final eruptive
phase (Phase 3b) extends from approximately August
1948 to the end of the eruption in 1952 and is dominated
by the extrusion of compositionally similar andesitic lavas
(60 wt % SiO2; 87Sr/86Sr 0·7042; McBirney et al., 1987).
Crustal material, either as xenoliths or xenocrysts, is very
rare in the Phase 3 eruptive material.
S A M P L E D E TA I L S A N D
A N A LY T I C A L M E T H O D S
Sample selection for melt inclusion study
For melt inclusion studies, tephra samples are often easier
to process because of the presence of naturally glassy melt
inclusions that can be analyzed without any heating and
rehomogenization (see below). Luhr (2001) analyzed
glassy melt inclusions in tephra samples from Paricutin,
but he was not able to cover the full compositional or temporal range observed in the erupted lavas because of the
restriction to tephra samples from the collections at the
Smithsonian National Museum of Natural History.
Samples for the present study were selected from a more
representative suite of lavas of known eruption age that
were the focus of previous studies (from the Smithsonian
collections; Wilcox, 1954; McBirney et al., 1987), thus providing a thorough chronological, geochemical and petrological context as the basis for the detailed melt inclusion
study. Selected lava samples were specifically chosen to
cover all of the eruptive phases as previously defined (see
details in Table 1), although this meant that any melt inclusions were likely to be crystallized and would have to be
rehomogenized to a glass prior to microbeam analysis.
Olivine is the dominant phenocryst phase in all lava
samples erupted prior to 1947. The early 1943 lava
(116293-7) contains 5 vol. % olivine phenocrysts up to
1·5 mm in length, with rare plagioclase microphenocrysts, and the groundmass is dominated by plagioclase4orthopyroxene. McBirney et al. (1987) noted that
plagioclase is present in lavas from 1943 to 1944, but it appears as a microphenocryst only after mid-1944, consistent
with our observations. Lava samples 116295-23 and
116289-8 contain varying amounts (generally less than
4 vol. %) of euhedral olivine phenocrysts up to 1mm in
length, and a groundmass dominated by plagioclase with
minor orthopyroxene and olivine. Lava samples from
early 1947 (116289-9) to late 1947 (116289-12) have decreasing olivine abundances, from 5 vol. % to 1^2 vol. %,
with a groundmass of plagioclase with minor olivine.
Orthopyroxene is the predominant phenocryst phase
after 1947 (2^3 vol. %; up to 0·9 mm), as the host
magma changed to an andesitic composition, although
experimental data demonstrate that olivine is a potential
stable phenocryst phase at low pressure (51 kbar;
Eggler, 1972). Euhedral olivine is present as a phenocryst
phase in Phase 3 lavas, but it is rimmed with orthopyroxene (up to 10 mm thick). Luhr (2001) argued that such rims
of orthopyroxene must have formed after eruption while
the lavas slowly cooled, as these rims are not observed on
olivine grains in rapidly quenched tephra. A few spinel
crystals were observed in the early lavas but are volumetrically insignificant (McBirney et al., 1987; Bannister et al.,
1998).
2189
JOURNAL OF PETROLOGY
VOLUME 52
NUMBER 11
NOVEMBER 2011
Table 1: Paricutin lava and xenolith samples
Smithsonian
Collection
Material
Eruption
Eruption
Melt
Rehomogenization
no. (NMNH)
no.
type
date
stagey
inclusionsz
temp. (8C)
116293-7
51-W-18
Lava
Feb. 1943
1
16
1158
108081
108081
Lava
Jan. 8, 1944
2
–
–
116295-27
W-47-27
Lava
Oct. 1944
2
–
–
116295-23
W-47-23
Lava
Sept. 1945
2
17
1175
116289-8
W-46-27
Lava
Sept. 18, 1946
2
12
1169
116289-9
W-47-9
Lava
Apr. 9, 1947
3a
10
1172
116289-12
W-47-30
Lava
Nov. 30, 1947
3a
14
1117
116289-13
W-48-5
Lava
Aug. 1948
3b
–
–
116289-15
FP-20-49
Lava
Dec. 13, 1949
3b
25
1112
116289-16
FP-20-50
Lava
Sept. 1, 1950
3b
–
–
116289-19
FP-16-52
Lava
Feb. 25, 1952
3b
32
1113
108126
108126
Xenolith
Unknown
Unknown
–
–
116289-20
51-W-1
Xenolith
May 1943
1
–
–
116293-4
51-W-6
Xenolith
1943
1–2
–
–
116293-5
51-W-7
Xenolith
1944
2
–
–
116289-23
51-W-8
Xenolith
1944
2
–
–
*Sample identification and eruption dates from the Smithsonian National Museum of Natural History database (NMNH
sample prefix removed).
yEruption stages follow those defined in the text.
zNumber of melt inclusions analyzed for major elements.
Whole-rock lava and crustal xenolith
analyses
Details of the studied lava and crustal xenolith samples are
summarized in Table 1. Whole-rock trace element compositions were analyzed by inductively coupled plasma mass
spectrometry (ICP-MS) at Washington State University,
using the methods described by Knaack et al. (1994), on aliquots of powder from 11 lava and five crustal xenoliths on
loan from the Smithsonian National Museum of Natural
History. The new trace element data are presented in
Table 2, together with accompanying whole-rock major
element data for all the samples from McBirney et al.
(1987).
Sample preparation for melt
inclusion study
Paricutin lava samples were hand crushed and sieved, and
olivine, pyroxene, and plagioclase grains were handpicked
from the 4250 mm fragments under a binocular microscope. Microscope observations showed that the melt inclusions were mostly to completely crystalline and
therefore had to be rehomogenized to a glass prior to analysis. Mineral phases in the melt inclusions could not be
identified petrographically prior to rehomogenization.
Groundmass-free mineral grains were reheated in a 1atm
Deltech vertical tube furnace at the University of Iowa.
Furnace temperatures were estimated from whole-rock
compositions, based on liquidus calculations from both
COMAGMAT (Ariskin et al., 1993) and MELTS
(Ghiorso & Sack, 1995; Asimow & Ghiorso, 1998). Oxygen
fugacity in the sealed tube was maintained slightly below
the QFM (quartz^fayalite^magnetite) oxygen buffer with
a CO2^H2 gas mixture. Total time of heating above
10008C was kept to 15 min, with 10 min at run temperature, based on the methods and rationale discussed by
Rowe et al. (2006, 2007). Samples were then rapidly
quenched, which resulted in glassy melt inclusions (Fig. 2).
Quenched grains were separately mounted and polished
to expose the melt inclusions. We examined melt inclusions
in both olivine and orthopyroxene, but we were unable to
recover inclusions for analysis from plagioclase crystals.
Electron microprobe and secondary ion
mass spectrometry analytical methods
Major element compositions of melt inclusions (including
S and Cl), host olivine and orthopyroxene grains, and
xenolith glasses, were measured by electron microprobe
analysis (EMPA) at Oregon State University on a
Cameca SX100 instrument. Detailed analytical methods
for analysis of melt inclusions and olivine grains have
2190
ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
Table 2: Major and trace element compositions of Paricutin lavas and xenoliths
Sample (NMNH):
116293-7
108081
116295-27
116295-23
116289-8
116289-9
116289-12
116289-13
Major elements (McBirney et al., 1987; wt %)
SiO2
54·59
55·39
55·71
55·79
56·13
57·05
58·39
TiO2
0·99
0·94
1·01
0·90
1·02
0·89
0·86
59·09
0·78
Al2O3
17·83
17·64
17·24
17·48
17·34
17·27
17·78
17·55
Fe2O3
2·01
2·16
2·06
1·83
1·74
1·42
1·87
2·04
FeO
5·43
5·46
5·48
5·30
5·42
5·21
4·51
4·27
MnO
0·12
0·13
0·13
0·12
0·12
0·12
0·12
0·11
MgO
5·44
5·43
5·61
5·75
5·58
5·64
4·03
4·03
CaO
7·25
7·18
6·98
6·81
6·99
6·94
6·75
6·46
Na2O
3·95
3·98
3·99
3·81
3·79
3·71
3·86
3·92
K2O
0·91
1·15
1·18
1·19
1·30
1·23
1·30
1·50
P2O5
0·27
0·35
0·33
0·30
0·36
0·29
0·30
0·08
H2Oþ
0·16
0·09
0·20
0·20
0·20
0·17
0·11
0·03
H2O
0·04
0·05
0·06
0·10
0·06
0·02
0·01
0·30
Total
98·99
99·95
99·98
99·58
100·05
99·96
99·89
100·16
Ni (ppm)
116
103
126
127
122
126
71
73
Cr (ppm)
144·7
176
155·2
170
152·7
162·4
78·5
88
Trace elements (Washington State University, ICP-MS, ppm)
La
13·26
17·70
18·88
18·01
18·57
16·64
18·00
19·46
Ce
28·58
37·47
39·72
37·57
38·79
34·49
37·30
40·17
Pr
3·86
4·93
5·16
4·89
5·05
4·48
4·89
5·21
Nd
16·61
20·39
21·30
20·09
20·75
18·42
20·08
21·10
Sm
3·82
4·54
4·68
4·33
4·58
4·07
4·32
4·63
Eu
1·30
1·52
1·52
1·42
1·47
1·31
1·37
1·41
Gd
3·71
4·31
4·44
4·12
4·30
3·83
4·00
4·10
Tb
0·59
0·68
0·70
0·64
0·67
0·58
0·62
0·63
Dy
3·45
3·99
4·07
3·82
4·01
3·41
3·61
3·75
Ho
0·70
0·79
0·81
0·76
0·79
0·69
0·71
0·75
Er
1·87
2·13
2·18
2·03
2·09
1·79
1·90
1·98
Tm
0·27
0·30
0·31
0·29
0·30
0·25
0·27
0·28
Yb
1·67
1·87
1·92
1·79
1·84
1·61
1·65
1·74
Lu
0·26
0·30
0·31
0·28
0·29
0·26
0·27
Ba
311
372
388
400
398
416
472
0·29
506
Th
1·01
1·57
1·72
1·72
1·81
1·65
1·64
1·82
Nb
5·25
8·87
9·62
8·30
9·09
7·08
6·90
7·19
Y
17·70
20·12
20·92
19·16
20·14
17·49
18·18
18·83
Hf
2·95
3·76
3·98
3·75
3·88
3·53
3·81
4·08
Ta
0·34
0·59
0·63
0·55
0·60
0·47
0·47
0·49
U
0·37
0·53
0·56
0·56
0·56
0·52
0·54
0·58
Pb
4·95
6·03
5·97
6·15
6·31
6·28
6·98
Rb
Cs
Sr
Sc
Zr
10·4
0·35
603
18·1
109
15·3
0·43
588
16·6
146
16·6
0·42
594
17·7
157
17·0
0·44
582
17·2
146
17·3
0·45
576
18·0
152
19·0
0·53
565
17·3
135
20·6
0·60
562
14·5
145
7·45
22·4
0·58
547
14·2
155
(continued)
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JOURNAL OF PETROLOGY
VOLUME 52
NUMBER 11
NOVEMBER 2011
Table 2: Continued
Sample (NMNH):
116289-15
116289-16
116289-19
108126
116289-20
116293-4
116293-5
116289-23
Av. xeno
Major elements (McBirney et al., 1987; wt %)
SiO2
59·77
60·24
60·07
71·93
70·88
72·61
71·00
75·95
TiO2
0·83
0·80
0·81
0·23
0·36
0·17
0·18
0·04
0·20
Al2O3
17·29
17·28
14·98
14·27
14·98
14·83
13·51
14·51
Fe2O3
1·21
1·19
1·37
0·80
1·52
0·55
0·64
0·25
0·75
FeO
4·95
4·59
4·39
1·25
1·53
1·51
1·43
0·27
1·20
MnO
0·11
0·10
0·10
0·06
0·05
0·06
0·06
0·03
0·05
MgO
3·72
3·55
3·73
0·55
1·17
0·32
0·53
0·05
0·52
CaO
6·28
6·14
6·16
2·79
1·65
2·96
3·13
1·05
2·32
Na2O
3·74
4·01
4·00
4·67
4·18
4·75
4·13
3·90
4·33
K2O
1·67
1·66
1·67
2·15
3·64
1·63
2·38
4·74
2·91
P2O5
0·12
0·04
0·03
0·19
0·11
0·39
0·47
0·13
0·14
H2Oþ
0·00
0·04
0·05
0·03
0·05
0·07
0·11
0·01
0·26
H2O
0·31
0·29
0·28
0·27
0·08
0·16
0·19
0·02
0·05
99·95
99·94
99·9
99·49
100·16
99·08
99·95
99·72
11
10
Total
100
17·3
Ni (ppm)
57
44
63
Cr (ppm)
67·8
66·5
76
4·7
18
72·47
15
18
15
37
17
Trace elements (Washington State University, ICP-MS, ppm)
La
20·26
20·32
20·24
14·13
21·94
14·65
13·74
5·87
14·06
Ce
41·48
41·60
41·05
26·11
41·65
26·56
25·19
13·61
26·62
Pr
5·33
5·32
5·19
2·95
4·70
2·97
2·88
1·97
3·10
Nd
21·46
21·37
20·86
10·30
16·46
10·46
9·98
8·97
11·23
Sm
4·54
4·51
4·37
1·85
3·30
1·82
1·78
2·90
2·33
Eu
1·39
1·37
1·31
0·54
0·68
0·53
0·52
0·18
0·49
Gd
4·16
4·09
3·90
1·49
3·08
1·51
1·48
3·26
2·16
Tb
0·64
0·63
0·60
0·23
0·50
0·23
0·23
0·59
0·36
Dy
3·72
3·71
3·56
1·37
3·09
1·37
1·34
3·75
2·18
Ho
0·74
0·73
0·69
0·27
0·62
0·28
0·27
0·77
0·44
Er
1·93
1·94
1·88
0·77
1·75
0·75
0·74
2·11
1·22
Tm
0·28
0·28
0·27
0·12
0·27
0·12
0·11
0·32
0·19
Yb
1·71
1·74
1·64
0·80
1·73
0·78
0·77
2·01
1·22
Lu
Ba
0·27
540
0·28
552
0·26
554
0·14
327
0·28
498
0·14
339
0·13
676
0·31
124
0·20
393
Th
1·96
2·02
2·14
1·80
21·32
1·76
1·71
8·26
6·97
Nb
7·51
7·53
8·00
3·87
5·14
3·98
3·76
3·85
4·12
Y
18·79
18·75
17·84
7·63
17·08
7·70
7·58
23·16
12·63
Hf
4·25
4·22
4·11
3·00
4·75
2·96
2·86
2·73
3·26
Ta
0·51
0·52
0·53
0·34
0·56
0·34
0·32
0·43
0·40
U
0·62
0·63
0·67
0·45
4·93
0·46
0·44
3·25
1·91
Pb
7·95
7·96
8·06
5·85
7·04
7·58
8·82
26·18
11·09
Rb
Cs
Sr
Sc
Zr
24·7
0·64
531
14·5
161
26·0
0·67
528
15·0
161
27·2
0·72
534
14·4
158
44·8
1·17
395
2·6
106
2192
157·3
2·30
50
8·5
153
34·0
1·14
441
2·9
107
62·2
1·23
458
2·8
102
149·9
2·82
52
1·9
45
89·7
1·73
279
3·72
102
ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
Fig. 2. Reflected light images of Type I and Type II melt
inclusions. Numbering represents the SiO2 content of melt inclusions
and the adjacent olivine forsterite composition. (Note the presence of
Type I and Type II inclusions in compositionally similar olivine.)
been provided by Rowe et al. (2011). Orthopyroxene grains
were analyzed using the same procedure and calibration
as olivine. Rhyolitic glasses were analyzed using the same
beam conditions as basaltic glass (7 mm spot, 30 nA
beam current, 15 keV accelerating voltage) with a linear
correction applied to recalculate Na and Si counts to a
zero time intercept. Rhyolitic glass standard (USNM
72854 VG-568) was repeatedly analyzed as an
intra-laboratory standard. KE-12 Obsidian Glass (Devine
et al., 1995) and Macusani Glass (Pichavant et al., 1987)
were also analyzed as secondary rhyolite standards. Major
element abundances, accuracy, and precision of basaltic
and rhyolitic glass standards are presented in Table 3.
Melt inclusion and xenolith glass trace element abundances were determined by secondary ion mass spectrometry (SIMS) at Arizona State University. Melt inclusion
trace element concentrations (88Sr, 89Y, 90Zr, 93Nb, 138Ba,
139
La, 140Ce, 144Nd, 147Sm, 151Eu, 158Gd, 162Dy, 174Yb) were
analyzed on a Cameca 3f ion microprobe, together with
30
Si and 42Ca for calibration and to allow correction for
potential olivine overlap. For trace element analysis we
used a 16O primary beam (0·8^1·2 nA) focused to
20^25 mm in diameter. Positive secondary ions were
accelerated to 4·5 keV and, following conventional energy
filtering techniques (Shimizu et al., 1978), ions with a
75 20 eV excess kinetic energy were allowed into the
mass spectrometer. Trace elements were analyzed in two
blocks (masses 88Sr to 139La and 140Ce to 174Yb) with 30Si
measured before and after each block for normalization.
Count times were 10 s (30Si), 20 s (42Ca, 88Sr), 30 s (masses
from 89Y to 144Nd, plus 158Gd and 162Dy) or 40 s (147Sm,
151
Eu, 166Er, 174Yb). Measured ratios (Mþ/30Si) were corrected for interfering oxides using rare earth element
(REE) oxide production values (MOþ/Mþ) from Zinner
& Crozaz (1986) and a 135Ba16O/135Ba oxide production
ratio of 0·054 (R. Hervig, personal communication). For
small inclusions, CaO concentrations were calculated
from 42Ca/30Si ratios and compared with CaO concentrations determined by EMPA. Where analyses are
determined to have overlapped onto the olivine or orthopyroxene host (lower calculated CaO wt % relative to
EMPA concentrations), concentrations were corrected
assuming an essentially linear dilution of the melt composition. Basalt glass BHVO-2 G was used as a calibration
standard, and accuracy and precision were based on
repeat analysis of BCR-2G run as an unknown. Precision
is generally better than 5% for masses lighter than 144Nd
and 6^9% for masses from 147Sm to 174Yb. Accuracy,
relative to preferred values for BCR-2G (GEOREM:
http://georem.mpch-mainz.gwdg.de), is better than precision for all elements except 158Gd (þ11%) and 174Yb
(13%). Additional details for accuracy and precision for
basaltic trace element analysis by SIMS are presented in
Table 4.
2193
JOURNAL OF PETROLOGY
VOLUME 52
NUMBER 11
NOVEMBER 2011
Table 3: Repeat analysis of secondary electron microprobe glass standards, with calculated precision and accuracy, run with
analysis of melt inclusions and crustal xenolith glasses
Sample:
BHVO-2G
BCR-2G
Lo-02-04ii
KE-12 Obsidianz
Macusani Glassz
Av.
Prec.
Acc.
Av.
Prec.
Acc.
Av.
Prec.
Acc.
Av.
Prec.
Acc.
Av.
Prec.
Acc.
(32)
(%)*
(%)y
(28)
(%)
(%)
(12)
(%)
(%)
(5)
(%)*
(%)y
(5)
(%)*
(%)y
Major elements (electron microprobe analysis, wt %)
SiO2
49·68
0·7
0·8
54·26
1·1
1·2
48·04
0·7
2·6
70·20
0·3
0·1
71·85
0·4
0·6
TiO2
2·76
1·3
1·1
2·33
1·0
2·1
2·44
1·7
8·2
0·30
14·0
8·6
0·05
71·9
12·6
Al2O3
13·68
0·5
0·6
13·87
0·3
1·2
12·37
0·9
1·6
8·03
0·4
6·0
16·15
0·5
1·9
FeO*
10·85
0·8
4·2
12·57
1·0
1·3
10·85
1·2
0·0
8·78
11·6
4·8
0·55
3·7
9·2
MnO
0·16
10·0
3·2
0·21
11·7
25·5
0·16
18·8
2·1
0·27
8·1
4·8
0·06
23·7
2·0
MgO
7·25
0·5
1·6
3·67
0·7
1·1
9·04
1·6
0·1
0·02
21·7
8·5
0·02
4·6
24·5
CaO
11·52
1·1
1·0
7·37
0·9
3·1
11·02
1·8
4·5
0·36
2·9
3·0
0·22
5·2
6·0
Na2O
2·14
2·2
12·2
2·88
3·0
7·6
2·31
2·9
7·5
7·28
1·5
0·0
4·22
1·3
2·2
K2O
0·50
4·3
1·6
1·77
2·0
1·2
0·53
4·1
11·9
–
–
–
–
–
–
P2O5
0·28
5·8
3·7
0·37
4·8
8·1
0·28
3·5
8·4
4·14
1·0
3·2
3·64
1·2
0·1
S
0·00
–
–
0·00
–
–
0·13
5·5
11·9
0·02
–
–
0·00
–
Cl
0·01
–
–
0·01
–
–
0·14
1·8
0·9
0·34
1·5
2·1
0·05
8·3
0·01
–
–
0·02
–
–
0·01
–
–
–
–
–
–
–
99·34
–
–
97·40
–
–
F
Total
98·90
99·82
–
15·4
–
96·84
*Precision (%) calculated as standard deviation/average 100.
yAccuracy (%) reported as the deviation from reported values. BCR-2G and BHVO-2G accepted values from GEOREM
(georem.mpch-mainz.gwdg.de/). LO-02-04ii is a natural glass (more variable major elements) with reported values from
Kent et al. (1999) and S and Cl from Rowe et al. (2006).
zReported values for accuracy calculations are from Pichavant et al. (1987) and Devine et al. (1995).
Trace element abundances in xenolith rhyolitic glasses
were determined on a Cameca 6f ion microprobe at
Arizona State University. In addition to elements analyzed
during melt inclusion analysis, 47Ti and 232Th were also
counted. We used a 16O primary beam (5 nA) focused to
20^25 mm in diameter. Positive secondary ions were
accelerated to 10 keV. Energy filtering similar to that
described for melt inclusion analysis was applied to the
rhyolite procedure. After an initial pre-sputter time of
180 s, ions with a mass less than 144Nd were counted for 1s
(93Nb, 139La, 144Nd counted for 4, 2, and 4 s, respectively),
and ions with a mass greater than 147Sm were counted for
5 s for each measurement cycle (25 cycles per analysis).
Measured Mþ/30Si ratios were corrected for interfering
oxides as described above. NIST 612 glass was used for calibration whereas NIST 610 glass was analyzed as an unknown before and after the analytical session. Precision of
rhyolitic glass trace element analysis (based on NIST 610
analyses; Table 4) is better than 5% for all elements
except Ba, Ce, Eu, and Th (510% precision), and accuracy
is better than 5% for all elements except Ti (12%),
Sr (7%), Nb (15%), Ba (10%), Ce (10%) and Dy (8%).
R E S U LT S
Melt inclusion screening and corrections
Care must be taken when interpreting either naturally
quenched or rehomogenized melt inclusion compositions.
Post-entrapment modification (predominantly host^melt
re-equilibration, and water loss) can dramatically alter
the composition of the inclusions (e.g. Danyushevksy
et al., 2000; Hauri, 2002). Corrections for these ‘natural’
processes and for the effects of rehomogenization can
result in significant modifications of the measured melt
compositions. In the following section we detail assumptions and corrections applied to melt compositions to
provide the clearest estimate of the trapped melt
compositions; we also provide both measured and corrected inclusion compositions as Supplementary Data
(available for downloading at http://www.petrology
.oxfordjournals.org/). Rehomogenization in a 1atm furnace requires the assumption of mineral^melt equilibrium
at the time of inclusion trapping. Melt inclusions in both
olivine and orthopyroxene were therefore recalculated to
be in Fe^Mg equilibrium with their respective hosts. A
constant Fe^Mg distribution coefficient (KD) of 0·30 is
2194
ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
Table 4: Repeat analysis of secondary ion mass spectrometry standards, with calculated precision and accuracy, run with
analysis of melt inclusions and crustal xenolith glasses
Sample:*
BHVO-2G
BCR-2G
NIST 610
Av.
Prec.
Acc.
Av.
Prec.
Acc.
Av.
Prec.
Acc.
(15)
(%)y
(%)z
(10)
(%)
(%)
(3)
(%)
(%)
Trace elements (secondary ion mass spectrometry; ppm)z
Ti
–
–
–
–
–
–
495·1
6·1
12·3
Sr
396·0
5·0
–
347·6
3·3
1·6
539·3
4·7
7·8
Y
26·0
6·3
–
34·9
4·1
0·2
462·1
0·1
2·6
Zr
170·0
3·6
–
184·4
1·8
0·2
426·9
4·1
3·0
Nb
18·3
3·6
–
12·8
3·6
2·3
498·6
2·1
15·9
10·9
Ba
131·0
3·9
–
686·7
1·9
0·5
476·2
10·1
La
15·2
5·1
–
24·4
3·3
1·3
157·1
2·3
0·1
Ce
37·6
4·3
–
51·5
3·4
3·6
500·2
7·4
10·5
Nd
24·5
3·1
–
27·8
4·1
4·0
431·9
4·6
0·3
Sm
6·1
8·5
–
6·4
8·0
3·4
456·2
0·4
1·2
Eu
2·1
10·8
–
1·9
24·7
1·4
467·8
6·6
1·4
Gd
6·2
13·2
–
7·5
8·7
10·8
436·1
1·9
3·8
Dy
5·3
11·8
–
6·5
6·4
0·4
449·3
1·7
5·1
Yb
2·0
12·4
–
3·0
8·7
13·0
466·3
3·1
1·0
Th
–
–
–
–
–
–
460·2
8·4
0·7
*BHVO-2G was used as a calibration standard with BCR-2G run as an unknown during melt inclusion analysis. NIST 612
was run as a secondary standard during analysis of the crustal xenolith glasses.
yPrecision (%) calculated as standard deviation/average 100.
zAccuracy (%) reported as the deviation from reported values with reported values from GEOREM (georem.mpch-mainz
.gwdg.de/; BHVO-2G and BCR-2G) and Pearce et al. (1997) (NIST 610).
typically assumed for olivine^melt equilibria (Roedder &
Emslie, 1970) and for basaltic melt inclusion corrections
(e.g. Rowe et al., 2009, 2011). Erlund et al. (2010) calculated
a KD value 0·34 0·02 based on olivine core compositions
and bulk tephra Mg-number using the method described
by Toplis (2005). However, a KD value of 0·32 0·01 best
fits the naturally quenched glass and olivine compositions
presented by Erlund et al. (2010) and for this reason we
have corrected the rehomogenized inclusion compositions
using an olivine^melt KDFe^Mg of 0·32. Orthopyroxene^
melt KDFe^Mg (0·245) is calculated after von Seckendorff
& O’Neill (1993) assuming an average orthopyroxene
Mg-number of 78. This Fe^Mg distribution coefficient is
similar to that derived by Roedder & Emslie (1970) of
0·23. Fe speciation is based on whole-rock ferric/ferrous
determinations by McBirney et al. (1987).
Melt^host re-equilibration (Fe loss; Danyushevsky et al.,
2000) in olivine-hosted melt inclusions is monitored by
comparing measured melt FeO concentrations with
observed KDFe^Mg, as this process should produce a negative correlation between these two parameters in a suite of
inclusions (Rowe et al., 2011). Low-Si melt inclusions (see
below) from samples 116289-15 and 116289-19 both display
evidence of host re-equilibration and were corrected using
the software provided by Danyushevsky et al. (2000). For
both of these samples, the final melt FeO* was chosen to
be equal to the whole-rock FeO, with Fe^Mg Kd values as
described above and Fe3þ/Fetotal from the whole-rock analysis (McBirney et al., 1987). Although the basic choice of
an FeO concentration equivalent to the whole-rock FeO
may be an oversimplification, given the potential variability in melt compositions, excluding Fe and Mg, potential
errors induced by this procedure are relatively small. To illustrate this, the agreement between measured and corrected SiO2 and TiO2 concentrations are shown in Fig. 3.
Regardless of the correction technique applied, MgO and
FeO have the highest potential for error as these components are heavily leveraged by the host mineral compositions. If olivine-hosted melt inclusions are instead simply
corrected to be in equilibrium with their host olivine, incompatible elements (and SiO2) vary by less than 5%,
whereas MgO and FeO vary by up to 30%. Only the
2195
JOURNAL OF PETROLOGY
VOLUME 52
NUMBER 11
NOVEMBER 2011
Corrected major element compositions of
melt inclusions
Fig. 3. Measured vs corrected inclusion compositions for TiO2 and
SiO2.
corrected melt inclusion compositions (Table 5) are discussed in the following sections and figures; the measured
melt compositions are provided as Supplementary Data.
Sulfur provides an excellent monitor for melt inclusion
leakage as it is quickly lost by degassing under low or atmospheric pressure. Melt inclusions with sulfur concentrations below the sulfur detection limit (70 ppm S) have
been removed from the present study as these potentially
represent breached inclusions that have completely
degassed during either eruption or the rehomogenization
process and therefore may have suffered secondary alteration or significant diffusion along fractures (Nielsen
et al., 1998). Although recent studies have demonstrated
the feasibility of diffusion or re-equilibration of melt
inclusions with the host melt through phenocrysts, this process is more difficult to monitor when inclusion trace
element compositions are not anomalous, and therefore
re-equilibration may be extremely minor (e.g. Spandler
et al., 2007). We therefore consider that trace and
minor element abundances in unbreached melt inclusions
reflect the actual compositions from the time of melt
entrapment.
Each of the sampled Paricutin lavas records two compositionally distinct populations of olivine-hosted melt inclusions (Figs 4 and 5; Table 5): one with lower SiO2
(Type I) and one with higher SiO2 (Type II), with a typical average gap of 4^8 wt % SiO2 between the two
types. As is demonstrated below, variations in sulfur content also serve as a distinguishing characteristic of the two
groups (Fig. 5). The compositional gap is present in both
the corrected and uncorrected datasets, indicating this is
not an artifact of the correction process. There is no relationship between the size of the melt inclusions and their
compositional type (Type I, 5^40 mm; Type II, 5^45 mm;
Fig. 5), and the same olivine grain can host both Type I
and Type II inclusions, although on average the Type II
inclusion host compositions have a lower Mg-number
(Fig. 4; Table 5). The absence of a correlation between inclusion size and composition and the fact that both inclusion types can be found in the same grain reduces the
likelihood that compositional differences are an artifact of
re-equilibration and diffusion (e.g. Spandler et al., 2007).
Similarly, olivine^melt re-equilibration cannot produce
the observed variations in the melt compositions, in particular the bimodal populations, based on the absence of a
compositional correlation to inclusion size (Danyushevsky
et al., 2002) and the requirement of a multi-phase crystallizing assemblage (see below). A similar bimodal population
was documented from a small volume basaltic eruption at
Dotsero Volcano, Colorado, and was interpreted to record
late-stage crustal assimilation and episodic crystallization
(Rowe et al., 2011).
Average Type I melt inclusions have major element compositions that are equivalent to or slightly elevated relative
to whole-rock compositions, with the noted exception of
SiO2 and Al2O3. Type I inclusions have SiO2 concentrations equivalent to whole-rock SiO2 concentrations in
Phase 1 lavas, but SiO2 is lower relative to whole-rock compositions by up to 4 wt % in eruptive Phases 2 and 3
(Fig. 4). Al2O3 and CaO concentrations are proportionally
enriched in all inclusions relative to whole-rock abundances, from 0·3 to 2·1 wt % and from 0·3 to 0·7 wt %, respectively (Fig. 6). Generally, TiO2 and P2O5 contents in
the melt inclusions decrease over the course of the eruption, similar to the whole-rock trends (Fig. 6). Al2O3 concentrations do not vary systematically and remain
relatively constant throughout Phase 3. K2O and SiO2 are
the only major elements to increase in the Type I inclusions
over the course of the eruption, with average K2O and
SiO2 in inclusions varying in the range 1·2^2·0 wt % and
54·2^57·6 wt %, respectively, from the beginning to
the end of the eruption. SiO2 concentrations in Type I inclusions are lowest in Phase 2 lavas with average SiO2 as
low as 52·5 wt %. From Phase 2 to Phase 3, SiO2 increases
2196
ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
by 4 wt %, similar to the increase in the whole-rock
compositions over the same time period but offset to systematically lower concentrations (Fig. 4; Tables 2 and 5:
McBirney et al., 1987).
Type II melt inclusion compositions are generally more
variable than Type I, both overall and within a given
sample. Average Type II melt inclusion compositions are
enriched in SiO2, TiO2, K2O, and P2O5, relative to both
the whole-rock and Type I compositions. Phase 1 Type II
compositions are significantly enriched relative to Phase 2
(e.g. 2·2 wt % average TiO2 vs 1·2 wt % average TiO2 in
Phase 2). The average concentrations of both CaO and
Al2O3 increase through Phase 1 and 2 and then decrease
significantly in Phase 3 lavas (Table 5). However, at
all times average CaO and Al2O3 contents in Type II
inclusions are lower than those of the whole-rock and
Table 5: Average corrected melt inclusion major and trace element compositions
Sample:
116293-7
116293-7
116295-23
116295-23
Comment:
Type I
Type II
Type I
Type II
Type I
Host:
Olivine
Olivine
Olivine
Olivine
Olivine
Av.
SD
Av.
SD
Av.
SD
Av.
116289-8
SD
Av.
Corrected inclusion composition (wt %)
SiO2
54·22
1·09
61·93
0·88
52·71
2·95
56·52
1·24
TiO2
1·24
0·11
2·17
0·49
1·06
0·38
1·18
0·29
52·51
1·11
Al2O3
20·29
0·73
15·47
1·61
17·79
1·66
16·82
2·20
18·41
6·65
FeO
4·60
0·69
4·39
0·97
6·91
1·30
5·97
1·34
Fe2O3
1·97
0·23
1·89
0·34
2·41
0·48
2·14
0·38
2·08
MnO
0·11
0·02
0·12
0·04
0·14
0·04
0·14
0·02
0·14
MgO
3·48
0·53
2·77
0·70
6·14
1·35
4·87
0·92
5·68
CaO
7·67
0·77
5·03
0·58
7·18
0·72
6·53
0·72
7·32
Na2O
4·61
0·32
3·24
0·25
3·79
0·46
3·77
0·46
4·05
K2O
1·20
0·10
2·04
0·24
1·20
0·35
1·47
0·26
1·36
P2O5
0·39
0·04
0·79
0·41
0·34
0·17
0·40
0·11
0·37
S
0·046
0·008
0·013
0·00
0·069
0·021
0·020
0·01
0·070
Cl
0·116
0·015
0·099
0·03
0·084
0·021
0·055
0·03
0·093
F
0·02
0·00
0·03
0·02
0·02
0·01
0·02
0·01
0·02
100·00
0·00
100·00
0·00
100·00
0·00
100·00
0·00
100·00
Host Mg-no.
Total
80·80
2·26
77·53
3·39
83·20
0·68
81·73
1·56
81·80
X (host) wt %
4·46
1·46
3·80
2·26
0·26
2·82
0·96
2·47
Dilution
1·04
1·04
0·99
1·03
0·06
1·01
Trace elements (ppm)
Sr
Y
Zr
Nb
Ba
668
18·5
121
6·3
351
51
1·2
7
0·5
41
440
24·9
186
9·8
465
153
12·3
89
3·6
30
580
0·4
18·9
0·4
146
1·1
8·9
0·6
395
5
595
17·5
135
8·6
405
171
8·7
76
4·4
91
507
24·1
165
11·2
555
La
13·3
0·9
17·3
7·6
16·7
1·1
17·0
7·4
21·5
Ce
30·1
2·7
37·0
13·3
37·0
1·2
35·1
15·3
47·2
25·8
Nd
17·7
1·9
23·7
10·6
20·8
1·6
19·4
9·1
Sm
4·0
0·6
5·3
2·0
4·8
0·1
4·5
1·7
5·7
Eu
0·9
0·2
1·1
0·2
1·1
0·2
1·3
0·6
1·7
Gd
4·0
0·3
4·0
1·0
3·7
0·6
4·0
1·5
5·8
Dy
3·3
1·2
3·9
0·2
3·4
0·5
3·6
1·9
4·2
Yb
1·9
0·3
2·0
0·6
2·2
0·3
1·9
0·8
2·2
(continued)
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JOURNAL OF PETROLOGY
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NUMBER 11
NOVEMBER 2011
Table 5: Continued
Sample:
116289-8
116289-9
116289-9
116289-12
116289-12
Comment:
Type II
Type I
Type II
Type I
Type II
Type I
Host:
Olivine
Olivine
Olivine
Olivine
Olivine
Olivine*
SD
Av.
SD
Av.
SD
Av.
SD
Av.
116289-15
SD
Av.
SD
Corrected inclusion composition (wt %)
SiO2
1·92
53·06
0·43
58·69
0·70
54·33
1·26
62·23
1·35
56·74
0·69
TiO2
0·19
0·97
0·04
1·41
0·15
1·02
0·19
1·46
0·10
0·93
0·07
Al2O3
0·98
18·60
0·45
15·33
0·72
19·16
2·38
15·45
0·62
19·30
0·46
FeO
0·65
6·64
0·62
6·43
0·63
5·60
1·49
4·38
0·26
4·94
0·01
Fe2O3
0·26
1·79
0·16
1·78
0·13
2·29
0·33
1·78
0·10
1·21
0·00
MnO
0·04
0·12
0·04
0·12
0·03
0·12
0·04
0·10
0·02
0·14
0·04
MgO
0·52
6·03
0·47
4·95
0·27
4·06
1·00
3·09
0·19
3·26
0·10
CaO
0·44
7·42
0·30
5·42
0·34
7·14
0·80
4·62
0·50
6·98
0·19
Na2O
0·27
3·65
0·38
3·70
0·25
4·23
0·45
3·74
0·32
4·26
0·36
K2O
0·13
1·25
0·12
1·62
0·09
1·52
0·17
2·47
0·15
1·87
0·23
P2O5
0·06
0·31
0·02
0·42
0·06
0·36
0·05
0·52
0·06
0·38
0·09
S
0·01
0·056
0·028
0·026
0·01
0·055
0·014
0·021
0·01
0·043
0·010
Cl
0·02
0·064
0·036
0·069
0·03
0·101
0·014
0·122
0·02
0·088
0·009
F
0·01
0·02
0·00
0·02
0·00
0·02
0·00
0·03
0·00
0·02
0·01
Total
0·00
100·00
0·00
100·00
0·00
100·00
0·00
100·00
0·00
100·00
0·00
Host Mg-no.
0·99
83·47
0·97
81·15
0·66
80·23
0·39
79·66
0·51
78·63
0·55
X (host) wt %
1·41
1·23
1·79
0·15
1·04
0·03
3·38
0·25
0·78
–
–
Dilution
–
0·98
–
0·98
–
–
0·99
–
1·03
–
0·99
Trace elements (ppm)
Sr
Y
Zr
Nb
Ba
36
2·7
33
2·0
35
585
20·1
154
8·7
454
44
4·2
26
1·6
63
389
28·6
214
14·1
752
37
593
3·6
20·9
36
168
2·2
8·9
52
515
30
3·5
33
1·8
75
287
29·0
253
13·4
812
26
2·9
16
0·8
83
542
19·0
167
8·5
565
35
0·7
7·0
0·5
11
La
2·9
19·1
3·7
28·0
4·0
19·8
3·8
26·9
1·7
19·1
0·9
Ce
4·1
40·0
8·6
58·4
7·6
41·0
7·7
60·2
6·9
41·0
1·2
Nd
2·5
21·2
4·4
32·4
3·1
22·5
4·8
32·1
2·2
22·3
1·6
Sm
0·8
4·8
1·4
7·0
1·0
5·5
0·9
6·4
0·7
4·9
0·5
Eu
0·2
1·2
0·5
1·9
0·4
1·4
0·3
2·0
0·3
0·9
0·4
Gd
0·9
4·1
1·3
7·3
0·8
4·3
1·7
7·9
0·6
4·1
0·6
Dy
0·9
3·9
1·0
5·1
1·3
3·7
0·7
5·6
0·3
3·3
0·5
Yb
0·3
2·3
0·4
2·5
0·1
2·2
0·3
2·5
0·2
1·9
0·2
(continued)
Type I inclusions. Type II melt inclusions are compositionally similar to the matrix glass of the contemporaneous
tephra deposits (Luhr, 2001; Erlund et al., 2010; Fig. 6).
Although Luhr (2001) expressed doubts about the ability
of orthopyroxene to preserve undegassed melt inclusions,
given the strong cleavage of the mineral, we have successfully collected major and volatile element data from this
mineral phase, which is present only in the Phase 3 lavas.
Orthopyroxene-hosted melt inclusions from Phase 3 lavas
tend to be smaller (515 mm) than the olivine-hosted inclusions (Fig. 5), which limited our ability to obtain trace
element analyses on them. They are compositionally similar to the olivine-hosted inclusions and can be similarly
subdivided into two populations based on their SiO2
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ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
Table 5: Continued
Sample:
116289-15
116289-15
116289-15
116289-19
116289-19
116289-19
116289-19
Comment:
Type II
Type I
Type II
Type I
Type II
Type I
Type II
Host:
Olivine
Opx
Opx
Olivine*
Olivine
Opx
Opx
Av.
SD
Av.
SD
Av.
SD
Av.
SD
Av.
SD
Av.
SD
Av.
SD
Corrected inclusion composition (wt %)
SiO2
63·48
1·19
58·88
0·68
63·28
–
57·24
0·69
63·09
2·38
57·62
2·60
64·47
TiO2
1·92
0·15
0·84
0·03
1·48
–
0·99
0·14
1·33
0·26
0·85
0·12
1·34
–
–
Al2O3
15·27
0·96
17·37
0·39
13·62
–
19·39
0·68
16·07
1·62
16·89
1·08
14·38
–
FeO
5·17
0·07
5·70
0·64
5·92
–
4·39
0·02
4·26
0·54
6·77
1·47
4·66
–
Fe2O3
1·17
0·01
1·46
0·14
1·33
–
1·37
0·01
1·24
0·12
2·20
0·40
1·56
–
MnO
0·07
0·02
0·12
0·03
0·08
–
0·14
0·04
0·06
0·01
0·13
0·03
0·13
–
MgO
3·33
0·00
3·29
0·34
3·59
–
3·18
0·06
2·97
0·41
3·62
1·05
2·25
–
CaO
3·85
0·09
6·68
0·25
4·45
–
6·83
0·26
4·28
0·96
6·26
0·59
4·43
–
Na2O
2·47
0·39
3·58
0·24
2·83
–
4·18
0·23
3·20
0·82
3·54
0·33
3·31
–
K2O
2·54
0·02
1·60
0·14
2·58
–
1·96
0·24
2·89
0·39
1·70
0·40
2·86
–
P2O5
0·58
0·00
0·31
0·03
0·63
–
0·35
0·05
0·45
0·10
0·28
0·03
0·52
–
S
0·019
0·00
0·046
0·011
0·011
–
0·039
0·010
0·019
0·01
0·048
0·021
0·007
–
Cl
0·120
0·03
0·086
0·015
0·170
–
0·092
0·011
0·110
0·02
0·078
0·013
0·078
–
F
0·03
0·00
0·01
0·00
0·03
–
0·02
0·00
0·02
0·01
0·01
0·00
0·03
–
100·00
0·00
100·00
0·00
100·00
–
100·00
0·00
100·00
0·00
100·00
0·00
100·00
–
80·09
0·27
79·41
0·84
79·11
3·27
77·91
–
1·52
1·01
1·85
3·52
2·40
–
1·03
–
–
Total
Host Mg-no.
78·17
0·26
80·77
1·70
81·57
–
X (host) wt %
1·70
0·14
2·24
1·70
3·40
–
Dilution
1·00
1·04
0·99
1·04
1·01
1·02
Trace elements (ppm)
Sr
–
–
–
–
–
–
Y
–
–
–
–
–
–
523
Zr
–
–
–
–
–
–
Nb
–
–
–
–
–
–
Ba
–
–
–
–
–
–
La
–
–
–
–
–
–
19·2
1·3
28·6
Ce
–
–
–
–
–
–
41·2
2·8
58·7
Nd
–
–
–
–
–
–
21·6
1·4
28·2
Sm
–
–
–
–
–
–
4·2
0·5
Eu
–
–
–
–
–
–
1·0
Gd
–
–
–
–
–
–
Dy
–
–
–
–
–
Yb
–
–
–
–
–
17·9
163
9·0
32
0·1
6
23·6
209
4·9
73
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
–
5·3
–
–
–
–
10·7
–
–
–
–
6·2
–
–
–
–
6·1
1·0
–
–
–
–
0·1
1·5
0·6
–
–
–
–
3·5
0·5
5·7
3·0
–
–
–
–
–
3·3
0·7
4·9
1·5
–
–
–
–
–
1·7
0·2
2·0
0·4
–
–
–
–
1093
5·8
–
–
26
13·5
137
–
589
0·3
396
164
*Compositions corrected using the spreadsheet and procedure described by Danyushevksy et al. (2000).
contents, although the low-SiO2 Type I population is dominant (Fig. 4, Table 5). The only significant difference between orthopyroxene- and olivine-hosted melt inclusions
is that Type I opx-hosted inclusions have SiO2 concentrations more similar to those of the whole-rock and up to
2 wt % greater than those of the Type I olivine-hosted
inclusions. MgO concentrations are also comparable in
Type I opx-hosted and olivine-hosted inclusions (Table 5).
Higher FeO concentrations in Type I opx-hosted inclusions
(up to 1·5 wt %) relative to olivine-hosted inclusions may
be a function of the Fe-loss corrections rather than a
geologically significant difference. Type I opx-hosted inclusions more closely approximate the whole-rock compositions for Phase 3 lavas than the Type I olivine-hosted
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JOURNAL OF PETROLOGY
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NUMBER 11
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Fig. 4. Major and trace element compositions, and host Mg-number of Type I (open symbols) and Type II (filled symbols) melt inclusions vs
eruption date. Whole-rock major element data are from McBirney et al. (1987) (black symbols and continuous line) and Luhr (2001). Dashed
and dotted lines, respectively, represent the average compositions for each group of olivine-hosted and orthopyroxene-hosted melt inclusions.
Luhr (2001) melt inclusion compositions are plotted for comparison (open circles).
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ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
Fig. 5. (a) SiO2 vs S for melt inclusions, divided into the three main
phases and inclusion Types I and II. (b) Histogram of olivine- and
orthopyroxene-hosted melt inclusion sizes for Type I (low-Si) and
Type II (high-Si) inclusion populations.
inclusions. Type II opx-hosted inclusions are compositionally similar to Type II olivine-hosted inclusions in a given
sample.
Trace element compositions of lavas and
melt inclusions
New trace element data for 11 whole-rock lava samples are
reported in Table 2. The trace element characteristics of
the Paricutin samples are summarized on a primitivemantle-normalized diagram (Fig. 7). They show
Fig. 6. TiO2, P2O5 and Al2O3 vs SiO2 in melt inclusions and
whole-rock samples demonstrating the overall trend of Type I melt inclusion and whole-rock compositions toward xenolith (filled circles)
and xenolith glass (open circles) compositions. Type II melt inclusions
follow fractional crystallization trends (f.c.) similar to the Luhr
(2001) and Erlund et al. (2010) tephra glass compositions.
enrichments in the large ion lithophile elements (LILE;
Cs, Rb, Ba, U, K) and Th, with negative anomalies at
Nb, Ta, P, and Ti; these features are typical for magmas
derived from subduction-modified mantle (e.g. Pearce &
Peate, 1995). The REE patterns show strong light REE
(LREE) enrichments (La/SmN ¼ 2·63 0·23), elevated
middle REE (MREE) to heavy REE (HREE) levels
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Fig. 7. Primitive mantle normalized whole-rock trace element patterns of lavas (black lines) and crustal xenoliths (gray lines). Primitive mantle
composition from McDonough & Sun (1995). Data from this study.
(Dy/YbN ¼1·43 0·03) and no negative Eu anomalies
(Eu/Eu* ¼ 1·01 0·03). The Phase 2 samples are distinctive
in having lower Zr/Nb (16^18) compared with those from
Phase 1 and Phase 3, a feature that is also apparent in
other published datasets (Luhr, 2001; Cebria¤ et al., 2011).
Large ion lithophile elements (Rb, Cs, Ba, Pb, K) show a
progressive increase and Sr shows a progressive decrease
with time from Phase 1 to Phase 2 to Phase 3 samples
(Fig. 4). Zr and the LREE show similar abundances in
the Phase 2 and Phase 3 samples that are higher than in
the Phase 1 samples, whereas Y and the HREE show
similar abundances in the Phase 1 and Phase 3 samples
that are lower than in the Phase 2 samples. Nb contents
increase from the Phase 1 to Phase 2 whole-rock
samples and decrease to Phase 3. In general, most incompatible trace elements show positive linear trends against
SiO2, with the Phase 2 samples displaced to higher
contents.
For the melt inclusions, only the olivine-hosted inclusions were large enough (greater than 20 mm) for trace
element analysis. Type I melt inclusions generally have
trace element abundances similar to whole-rock compositions (Fig. 4; Tables 2 and 5). Sr concentrations in the
melt inclusions exhibit the greatest deviation from the
whole-rock compositions, particularly the Phase 1 and
Phase 3b samples, with concentrations decreasing over
the course of the eruption (Fig. 4). Trace element abundances in Type II melt inclusions have a significantly
greater deviation from the whole-rock compositions than
the Type I inclusions, particularly in Phase 3 lavas, with
concentrations up to double the whole-rock values (e.g.
Ba, Fig. 4). Only Sr concentrations in Type II melt inclusions are substantially below the whole-rock and Type I inclusion values, with abundances 50% lower than in the
whole-rock lavas.
Volatile compositions of melt inclusions
Sulfur concentrations in undegassed Type I melt inclusions
are variable with average abundances from 400 to
700 ppm and the highest concentrations in Phase 2 and
Phase 3a samples (Fig. 5; Table 5). Phase 1, Type I melt inclusions have sulfur concentrations less than 600 ppm (in
Fo83^76 olivine), significantly lower than observed in Phase
1 tephras by Luhr (2001; 1000^1200 ppm S in Fo84 olivine)
and Johnson et al. (2009; 1320^1750 ppm S in Fo85^87 olivine). Type II melt inclusions have consistently lower sulfur
abundances (average concentrations from 100 to 300 ppm
S), relative to Type I inclusions, although concentrations
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ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
are greater than the matrix glass S concentrations in the
tephra samples (40 ppm; Luhr, 2001).
Average chlorine abundances vary from 500 to 1200 ppm
Cl (Table 5) with no systematic differences between Type I
and Type II inclusions. With the exception of the Phase 1
samples that have anomalously high Cl (average
1150 ppm), average Cl concentrations in Type I inclusions
remain relatively constant throughout the eruption and
display no systematic variations (averages range from 630
to 990 ppm; Fig. 8a). In contrast, average Cl/K ratios
Fig. 8. (a) Cl/K vs chlorine concentrations for Type I melt inclusions. (See Table 1 for clustering of samples into eruptive phases.) Bulk (BC),
Lower (LC), Upper (UC) and Middle (MC) crust compositions from Rudnick & Gao (2004). Decreasing Cl and Cl/K represents a degassing
path, whereas constant Cl and decreasing Cl/K (increasing K) represents crustal contamination. (b) Cl/K vs K2O/TiO2, for Type I
olivine-hosted melt inclusions. (Note the anomalous Cl/K ratios of melt inclusions in the Phase 1 lava.)
2203
JOURNAL OF PETROLOGY
VOLUME 52
decrease over the course of the eruption from 0·12 to 0·06
in Type I inclusions (Fig. 8). The negative correlation between K2O/TiO2 and Cl/K despite relatively constant Cl
(excluding inclusions which record degassing as illustrated
in Fig. 8) indicates that Cl/K variations in undegassed
melt inclusions are being driven largely by increasing
K2O concentrations as a result of the high K2O and
low Cl of likely crustal contaminants (Fig. 8). Chlorine
degassing as indicated in Fig. 8 is also supported by
decreasing sulfur contents in these melt inclusions (see
Supplementary Data). Cl/K is also negatively correlated
with other indices of crustal contamination, such as Ba/
Nb and SiO2 wt % in melt inclusions. Type II inclusions
have average Cl/K ranging from 0·08 to 0·03, but show significantly more internal variation.
Host mineral compositions
Olivine
Olivine host compositions were measured adjacent to each
melt inclusion, and their variations with eruption date are
shown in Fig. 4. Average olivine compositions hosting
Type I inclusions range from Fo80·0 to Fo83·5; Luhr (2001)
reported a similar range in olivine compositions adjacent
to melt inclusions in tephra samples from 1943 to 1948
(Fo80·0 to Fo84·4). Average olivine compositions hosting
Type II inclusions extend to more evolved compositions,
ranging from Fo82·0 to Fo77·5. Olivine compositions hosting
Type I and Type II inclusion populations have similar temporal variations. Average Phase 1 olivines (Type I ¼ Fo80·8,
Type II ¼ Fo77·5) are more evolved than later Phase 2 olivines (Type I ¼ Fo82·9, Type II ¼ Fo81·6). This result is in contrast to the olivine core compositions in basal tephra
deposits reported by Erlund et al. (2010), which indicate
that the early eruptive material was particularly mafic
with olivine compositions up to Fo88·4. We see no evidence
of such Mg-rich olivines in the early erupted lavas.
Average olivine compositions become progressively more
fayalitic through Phase 3a and into Phase 3b before
becoming more forsteritic at the end of the eruption
(Type I ¼ Fo80·9, Type II ¼ Fo79·4; Fig. 4). As previously
stated, both Type I and Type II inclusions may be found
within the same olivine grain. In these instances, with few
exceptions, the olivine host measured adjacent to the inclusions is not appreciably different between the two types.
Orthopyroxene
Orthopyroxene grains from Phase 3b lavas are more
variable, in terms of Mg-number [molar Mg/
(Mg þ FeT) ¼ 72^82], than the coexisting olivine grains
(Fig. 4). Despite the limited number of Type II melt inclusions found in orthopyroxenes, host compositions for both
Type I and Type II inclusions are indistinguishable. The
overall range of orthopyroxene compositions is constant
through eruptive Phase 3b; however, the average
Mg-number decreases from 81 to 79 from 1949 to 1952, in
NUMBER 11
NOVEMBER 2011
contrast to the increase in average Fo-number in olivine
grains over the same time interval (Fig. 4).
Crustal xenoliths
Crustal xenoliths are predominantly granitic in composition (70^76 wt % SiO2), with melting textures ranging
from veinlets cross-cutting intact granitic fragments to
completely pumiceous rhyolitic glass. Samples of xenoliths
cover the full textural range, thus allowing us to potentially examine both bulk and partial assimilation of crustal
xenoliths to better constrain the nature and timing of assimilation. Older trace element analyses of xenoliths were
obtained by a combination of techniques, including X-ray
fluorescence, atomic absorption and neutron activation
(McBirney et al., 1987). New trace element analyses by
ICP-MS provide a more internally consistent dataset
(Table 2). Bulk xenoliths are predominantly depleted in
high field strength elements (HFSE; Nb, Ta, Zr, Hf, Ti)
and enriched in LILE (e.g. Cs, Rb, K) relative to the
whole-rock lava compositions (Table 2, Fig. 7). The xenolith
compositions fall into two groups, with one group having
higher Al2O3 and Sr and lower K, Rb, Y, HREE, Th and
87
Sr/86Sr compared with the other group. Relative to the
basalts, the bulk xenoliths generally have slightly higher
Zr/Nb (12^30), and higher Ba/Nb (32^180), Th/Nb
(0·5^4·1) and Th/Yb (1·8^12·3). By comparison, lower crustal xenoliths from the Valle de Santiago Maar field in the
northern Michoacan^Guanajuato volcanic field have comparable Sr abundances but are depleted in LREE, Zr,
Rb, and Th and enriched in Yb relative to Paricutin xenoliths (Urrutia-Fucugauchi & Uribe-Cifuentes, 1999).
Glass from three crustal xenoliths with textures varying
from partially melted to frothy and highly vesicular
(Fig. 9) was analyzed to determine the variability in
major and trace element abundances (Table 6). Glass compositions are predominantly high-silica rhyolite (SiO2
71^79%), with only a few dacite analyses (SiO2 65^68%).
In each xenolith, the glass compositions have higher K2O,
lower Na2O and much lower Th contents than the bulk
composition, but otherwise exhibit similar compositional
features. For example, xenolith 116289-23 has higher K2O
and lower FeO, MgO, Sr, Ba and LREE contents
compared with the other two xenoliths; these differences
are also seen in the glass compositions. In general, the
dacitic glasses are more trace element enriched
compared with the rhyolitic glasses. Sr concentrations
range from 9 to 41000 ppm whereas Ba ranges from
31 to 691ppm. Ti is similarly variable, with abundances
from 14 to 9500 ppm (Table 6). The glasses show a significantly greater range in Ba/Nb (9^6700) than the bulk
xenoliths (32^180), but generally have higher values as a
result of their high Ba and very low Nb concentrations.
K2O/TiO2 ratios are similar in the glasses and bulk xenoliths, ranging from 5 to 15, except where TiO2 contents
are 50·1 wt %, when higher ratios (50^190) are observed.
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ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
Fig. 9. Photomicrographs of crustal xenoliths showing varying amounts of melting. Textures range from (a) nearly completely crystalline to
(b) glassy groundmass but intact grains to (c, d) mostly glassy and pumiceous. Black scale bar in each image represents 1·25 mm. Sample
116289-20 (a) did not contain any glass for analysis.
Rare earth element concentrations are generally lower in
the glasses compared with the bulk xenoliths, although
with significant variability and overlap of HREE
abundances.
DISCUSSION
In the following discussion, we integrate our new
whole-rock and melt inclusion data with previously published data to develop a consistent model for the temporal
and compositional development of the Paricutin magmatic
system. Earlier studies (e.g. Wilcox, 1954; McBirney et al.,
1987; Luhr, 2001), including our own, were based on the
sample suite that was collected as the Paricutin eruption
progressed, which is archived at the Smithsonian National
Museum of Natural History. Cebria¤ et al. (2011) recently
published new elemental and Sr^Nd isotope data on a
newly collected suite of lavas, whose eruptive chronology
was estimated from satellite photos and the detailed descriptions of the eruptive history. Other recent studies
(Pioli et al., 2008; Erlund et al., 2010) considered newly collected sample profiles through the tephra deposits,
primarily to investigate the physical volcanology of the
eruption and its relationship to the changing composition
of the erupted magma. The sample stratigraphy and compositions allowed the tephra to be correlated to the different eruptive phases but not to the detailed chronology
available for the Smithsonian sample collection.
In the first part of the discussion we explore the link between the Type II inclusions and the tephra matrix glasses.
The remainder of the discussion focuses on temporal variations in the composition of the Type I melt inclusions
and their host lava compositions, and how these relate to
processes at a deeper level in the magma plumbing system.
Origin of Type II high-Si melt inclusions
Type II melt inclusions, identified by their overall higher
SiO2 concentrations, typically record a greater compositional range than the Type I inclusions within a particular
whole-rock sample. They are displaced to higher TiO2,
P2O5 and K2O values and lower Al2O3 and CaO values
relative to the Type I inclusions and the whole-rock lava
and tephra samples, but overlap in composition with the
matrix glasses of the tephra samples (Fig. 6; Luhr, 2001;
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NUMBER 11
NOVEMBER 2011
Table 6: Representative major and trace element analyses of crustal xenolith glasses
Sample:
116289-23 116289-23 116289-23 116289-23 116293-5 116293-5 116293-5 116293-5 116293-4 116293-4 116293-4 116293-4 116293-4
Analysis: 23-1-2
23-1-4
23-1-7
23-1-8
5-1-1
5-1-5
5-1-7
5-1-9
4-1-1
4-1-5
4-1-6
4-1-9
4-1-11
Major elements (wt %)
SiO2
71·98
75·33
73·32
68·39
75·80
64·92
73·43
77·30
74·52
77·03
77·10
71·74
TiO2
0·00
0·07
0·00
0·05
0·04
0·00
0·07
0·03
0·42
0·00
0·00
0·25
64·13
1·14
Al2O3
16·04
13·61
14·82
18·32
13·74
21·70
14·84
12·83
12·98
12·95
12·93
14·31
13·94
FeO
0·19
0·34
0·13
0·13
1·13
0·89
1·60
1·18
2·69
1·23
1·12
3·57
7·60
MnO
0·00
0·00
0·00
0·00
0·02
0·00
0·03
0·00
0·14
0·01
0·04
0·18
0·29
MgO
0·07
0·09
0·03
0·03
0·26
0·17
0·41
0·24
0·80
0·45
0·35
0·99
2·24
CaO
0·48
0·38
0·32
0·82
0·54
4·24
0·77
0·58
2·29
1·59
1·50
1·90
2·75
Na2O
4·87
3·97
4·35
5·54
4·54
6·29
4·70
4·07
3·89
4·51
4·55
4·49
4·98
K2O
6·53
6·07
6·43
6·97
3·59
2·23
3·71
3·50
2·22
2·60
2·50
2·56
1·36
S
0·00
0·004
0·008
0·00
0·00
0·005
0·006
0·00
0·002
0·00
0·008
0·004
0·003
Cl
0·012
0·012
0·011
0·009
0·018
0·005
0·010
0·013
0·012
0·011
0·009
0·024
Total
100·18
99·88
99·43
100·23
99·67
100·45
99·58
99·74
99·96
100·39
100·11
100·02
0·015
98·54
Trace elements (ppm)
Ti
15·3
18·4
14·1
17·2
343
Sr
11·2
11·3
9·1
99·1
129
Y
6·3
2·9
1·8
5·7
5·1
13·2
Zr
2·4
0·1
1·1
0·1
31·2
1·7
0·1
253
223
1642
210
135
291
5·2
562
1·7
16·2
30·3
24·8
445
2·8
10·3
5·2
93·3
0·1
0·1
31·7
La
1·3
1·1
1·5
1·4
6·8
151
6·6
5·4
9·1
4·2
3·3
10·9
Ce
1·4
1·7
2·3
1·9
12·2
227
11·9
8·9
16·4
7·7
5·1
17·7
Nd
0·6
0·6
0·9
0·5
3·6
62·9
4·0
3·2
5·6
3·0
1·8
5·8
44·1
Sm
0·5
0·2
0·2
0·3
0·8
10·3
0·9
0·5
1·5
0·5
0·3
1·5
11·2
Eu
0·2
0·3
0·1
0·8
2·5
2·9
3·0
2·3
2·1
2·6
2·3
2·9
2·9
Gd
0·6
0·4
0·2
0·5
0·9
7·4
1·2
1·1
2·6
1·2
0·5
2·1
13·2
Dy
1·3
0·6
0·5
1·0
1·4
5·6
1·5
1·0
4·0
2·4
0·4
2·3
18·9
Yb
0·6
0·3
0·2
0·8
0·6
0·7
0·7
0·6
1·6
3·6
0·4
1·0
7·2
Th
0·2
0·3
0·2
0·2
0·7
0·2
0·7
0·6
0·2
0·1
0·4
0·4
0·4
Erlund et al., 2010). Despite the relatively high forsterite
content of the host olivine (Fo82·0^77·5) Erlund et al. (2010)
demonstrated that the olivine compositions were in equilibrium with the tephra matrix glass. Luhr (2001) noted
that tie-lines connecting bulk tephra samples with their
matrix glass compositions were different in orientation
(e.g. increasing K2O and TiO2) from the main temporal
trend through the bulk lava and tephra samples (e.g.
increasing K2O and decreasing TiO2). Luhr (2001) and
Erlund et al. (2010) showed that the compositions of each
tephra matrix glass could be derived from its bulk tephra
composition through significant fractional crystallization
(up to 40%) of plagioclase þ olivine (orthopyroxene
in the Phase 3 samples), phases that form the groundmass
of the Paricutin lavas. The constant K2O/TiO2 of each
428
576
0·1
81·0
1003
53·3
408
0·3
9497
226
0·3
691
7·4
370
1848
43·0
867
1·9
16·6
23·1
467
Nb
420
0·6
4·8
47·9
280
Ba
250
2·1
92·8
1016
670
10·7
510
75·2
676
46·2
101
bulk tephra^matrix glass pair indicate that titanomagnetite is not part of this crystallizing assemblage, consistent
with petrographic observations and experimental constraints (Eggler, 1972; McBirney et al., 1987). Similarly, the
constant P2O5/TiO2 ratios indicate an absence of apatite
crystallization.
Major element variations indicate that the Type II melt
inclusions were formed through a similar crystallization
process to the tephra matrix glasses, and the trace element
data are also consistent with a model dominated by fractional crystallization of plagioclase þ olivine orthopyroxene. Trace elements incompatible in these phases (e.g.
Rb, Ba, Y, Zr, Nb, REE) are all enriched in the Type II inclusions relative to the Type I inclusions and whole-rock
samples, whereas Sr contents (compatible in plagioclase)
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ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
are lower (Fig. 4), consistent with the lower CaO and
Al2O3 contents (Figs 4 and 6). This indicates that Type II
melt inclusions record a very late-stage crystallization
event, as feldspar is not a phenocryst phase in most of
these samples but is present in the groundmass.
Although the Type II inclusions typically record a large
compositional range, there is little systematic variation in
the inclusion compositions that would support progressive
assimilation in the evolution of Type II inclusions (e.g.
increasing indices of crustal contamination with melt evolution or host phenocryst composition). K2O/TiO2 ratios
are typically considered an excellent indicator of contamination in Paricutin lavas (e.g. McBirney et al., 1987; Luhr,
2001; Erlund et al., 2010), as both elements are incompatible
in the crystallizing assemblage, whereas likely crustal
assimilants have significantly higher K2O/TiO2 (e.g. typically 6^13 in the bulk crustal xenoliths and 10^8000 in
the xenolith glasses; Tables 2 and 6) compared with the
lavas (K2O/TiO2 0·8^2·1). Therefore, contaminated
magmas will have higher K2O/TiO2 values, and yet
K2O/TiO2 does not significantly increase with SiO2 in
Type II inclusions within a single sample (Fig. 10).
Although there is no direct correlation with SiO2, the
range in K2O/TiO2 values at constant SiO2 observed
in some Type II inclusions may record minor
assimilation; however, this variation may also simply be
the result of natural variation in very late-stage groundmass crystallization (see below). Trace element ratios such
as Ba/Nb that are higher in contaminated magmas are
likewise similar to whole-rock ratios for single samples, except for high Ba/Nb in Type II inclusions from
one of the Phase 3 samples (116289-19) toward the end
of the eruption, from which there is also significant
evidence for crustal assimilation from whole-rock
compositions. Indices of contamination therefore suggest
that no significant assimilation occurred during the
late-stage crystallization and evolution of the Type II melt
inclusions.
It was observed at the time of eruption that although
denser and relatively degassed lavas, including the lava
samples from this study, were erupted from a vent site at
the base of the cone, the crater remained the site of continued degassing and explosive eruptions (Krauskopf, 1948).
To explain this essentially simultaneous activity at two different vent sites, more recent models have suggested a shallow separation of volatiles at very shallow levels, either at
the base of the cone or just below, after which point the
more degassed magma erupts from the base of Paricutin
(Krauskopf, 1948; McBirney et al., 1987; Pioli et al., 2008).
This model is supported by the low sulfur concentrations
Fig. 10. K2O/TiO2 vs SiO2 for a representative sample of Phase 2 (116295-23) and of Phase 3b (116289-19) melt inclusions (Type I and Type II
inclusions).
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JOURNAL OF PETROLOGY
VOLUME 52
in Type II inclusions relative to Type I inclusions (Fig. 5),
potentially indicating low-pressure degassing either before
or during inclusion entrapment. Additionally, tephra samples are notably lacking in Type II melt inclusions, with
the exception of only a few inclusions in a single sample
(Luhr, 2001). This suggests that entrapment of the Type II
inclusions probably occurred at shallow levels, after separation of a gas-rich component in the main conduit. This is
supported by sulfur concentrations in Type II inclusions
above that of the tephra glasses reported by Luhr (2001),
again suggesting that crystallization of Type II inclusions
did in fact occur shortly before complete degassing and
syn- or post-eruptive cooling of the erupted lavas.
The presence of high-Si melt inclusions in relatively high
forsterite olivine may be explained if the inclusions were
trapped within embayments or necks in the olivine that
were sealed very late, such that the olivine and melt were
not in equilibrium. This would explain the similar host forsterite content between Type I and Type II inclusions, such
that inclusions are trapped predominantly in the higher
forsterite host (except where sealed off). Rowe et al. (2011)
identified a similar high-Si inclusion population from a
lava flow at Dotsero Volcano, Colorado. Textural evidence
indicated that the Dotsero high-Si population resulted
from closing off of embayments in olivine hosts, resulting in the juxtaposition of higher Si, fractionated melts
with relatively high forsterite olivine. Given the twodimensional nature of the polished melt inclusions, however, this is difficult to assess. High-Si melt inclusions in
Fig. 2 show no obvious signs of elongation that may support
this model and do not appear distinctive in form from
Type I inclusions. Rehomogenization of the melt inclusions, had they not been completely sealed, would have resulted in complete volatile loss (S, Cl), so it is likely that
these inclusions were completely enclosed, in contrast to
the high-Si inclusion population at Dotsero Volcano
(Rowe et al., 2011). Regardless of the method of entrapment,
the assumption of olivine^melt equilibrium in the correction process does not strongly affect these melt compositions (the correction is generally less than addition or
subtraction of 4 wt % olivine), as indicated by comparing
the measured and corrected melt compositions for this
population (see Supplementary Data), nor does it affect
our interpretations and conclusions.
Eruptive Phase 1, a compositionally
distinct initial magma batch
The earliest magmatism at Paricutin (eruptive Phase 1)
was only sparsely sampled during the course of the eruption, despite the fact that it makes up 15% of the total
erupted volume (McBirney et al., 1987); however, coverage
has been improved recently by extensive sampling of the
basal tephras (Pioli et al., 2008; Johnson et al., 2009;
Erlund et al., 2010). These data confirm that there is a distinct compositional break between the Phase 1 and Phase
NUMBER 11
NOVEMBER 2011
2 magmas, highlighted by a clear gap in K2O contents
(Phase 1 51 wt %; Phase 2 41 wt %; Luhr, 2001; Erlund
et al., 2010). Although there is compositional overlap for
other major elements, the whole-rock compositions of
Phase 1 magmas have comparable MgO and lower SiO2
compared with the Phase 2 lavas, whereas Phase 1 melt inclusions on average have higher SiO2 and lower MgO
compared with Phase 2 lavas (Fig. 4; Tables 2 and 5).
Luhr (2001) and Erlund et al. (2010) showed that fractional
crystallization crustal assimilation could not explain
these minor and major element differences, and argued
that Phases 1 and 2 were characterized by separate, compositionally distinct, magma batches.
Trace element data confirm this model of distinct
magma batches. Despite having more primitive compositions (higher MgO, higher Cr and Ni; McBirney et al.,
1987), Phase 2 lavas are enriched in nearly all incompatible
trace elements (excluding Sr, Ti, and Sc) relative to Phase
1 lavas (Table 2). Basal phase 1 tephras, identified by
Johnson et al. (2009) and Erlund et al. (2010), however,
have lower SiO2 and higher MgO concentrations and a
more depleted incompatible trace element signature
(excluding Sr, which is comparable with Phase 1 lavas)
than either Phase 1 or 2 lavas (Table 2; Johnson et al.,
2009). Fractional crystallization of olivine plagioclase
cannot produce the observed differences in ratios of
highly incompatible elements between Phase 1 (e.g. Zr/Nb
21) and Phase 2 (e.g. Zr/Nb 17) magmas (Fig. 11).
There is also no change in the calculated Eu anomaly between Phase 1 and Phase 2 lavas, further suggesting that
fractionation of plagioclase was not driving the changes in
trace element concentrations. However, within Phase 1,
melt inclusions (Type I and Type II) indicate concurrent
plagioclase and olivine fractionation, based on an increasing negative Eu anomaly and decreasing Sr concentrations
(although more scattered) with decreasing host forsterite
content. Trace element compositions of melt inclusions
from Phase 1 (February^July 1943) lavas also support a
model in which the Phase 1 lavas are derived from a different magma batch compared with the subsequent Phase 2
lavas. For ratios of highly incompatible elements, such as
Zr/Nb, there is no compositional overlap of Type I inclusions in magmas of the two eruptive phases: inclusions in
Phase 1 magmas have Zr/Nb of 18^20, whereas inclusions
in Phase 1 magmas have Zr/Nb of 15^17, consistent with
the observed difference in the whole-rock compositions.
Phase 1 magmas have similar Th/Nb whole-rock ratios to
the Phase 2 magmas, but higher Ba/Nb ratios (Phase 1
450, Phase 2 550); these differences cannot be explained
by assimilation of crustal material similar to the analyzed
xenoliths, which have elevated Ba/Nb (average of 95) and
Th/Nb relative to the lavas (Table 2). Further evidence
that shallow crustal assimilation is not the cause for the
distinct change in composition comes from the
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ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
Fig. 11. (a) Whole-rock variation of Zr/Nb vs K2O/TiO2. Included for comparison are bulk tephra analyses from Luhr (2001) and Cebria¤ et al.
(2011), incorporated into each eruptive phase based on eruption date. Continuous curves represent simple mixing lines between the average
xenolith composition and xenolith 116293-5 (Table 2) and a Phase 2 and an initial Phase 3 lava. The average xenolith (used in previous studies)
and 116293-5 (best fit to Phase 3 assimilation by a single xenolith) compositions are utilized to demonstrate potential assimilation in Phase 3
and to illustrate that assimilation by Phase 2 lavas cannot produce Phase 3 lavas. Tick marks are 5 wt % increments. The discrepancy between
Phase 1 and Phase 2 lavas should be noted.‘Assimilation’ and ‘mixing’ trends are schematic only, representing potential assimilation of average
xenoliths by Phase 1 lavas and mixing between Phase 2 and Phase 3a (discussed in the text). Melt inclusions follow similar trends but are
offset probably as a result of a calibration difference between the SIMS and ICP-MS and are therefore not plotted. (b) Ba/Nb vs K2O/TiO2
for whole-rock (large symbols) and melt inclusions (small symbols). Also plotted are the whole-rock analyses of Luhr (2001), Cebria¤ et al.
(2011), and the basal tephra melt inclusion analyses from Johnson et al. (2009).
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JOURNAL OF PETROLOGY
VOLUME 52
indistinguishable 87Sr/86Sr ratios of Phase 1 and Phase 2
lavas despite the elevated 87Sr/86Sr of the local crust
(McBirney et al., 1987).
Comparing melt inclusion and host crystal compositions
provides a means to evaluate the possible links between
contamination and fractionation during the evolution of
Phase 1 magmas using olivine forsterite content as an
index of fractionation. Indices of crustal assimilation (e.g.
Ba/Nb, K2O/TiO2, Sr/Nd) in olivine-hosted melt inclusions (Type I) show no systematic evidence of progressive
contamination with melt evolution. K2O/TiO2 and Cl/K
ratios remain constant with decreasing forsterite content
(from 83 to 78). It has previously been suggested that Cl/
K may provide a means to identify crustal contamination
(Rowe et al., 2009), given the relatively high K2O content
of average upper crustal rocks (Rudnick & Gao, 2004)
and the negative correlation between Cl/K and other indices of contamination, such as Ba/Nb, SiO2 or K2O/TiO2
in olivine-hosted melt inclusions (Fig. 8). Ba/Nb is more
variable in Phase 1 melt inclusions, ranging from 53 to 60.
However, these variations are non-systematic and may reflect minor heterogeneity of the melt.
Type I melt inclusions record significant differences in
volatile concentrations between Phase 1 and Phase 2 samples. Phase 1 magmas from this study have higher Cl
(1380^900 ppm) and lower S contents (360^590 ppm) compared with the subsequent Phase 2 magmas (Cl 290^
950 ppm; S 490^930 ppm). However, based on the melt inclusion data from Luhr (2001) and Johnson et al. (2009)
Phase 1 melt inclusions from tephra samples have higher S
than Phase 2 melt inclusions. Cl/K ratios of Type I inclusions in the Phase 1 and Phase 2 samples are also significantly different, with Cl/K in Phase 1 samples higher than
later erupted material as a result of both lower K2O and
higher Cl in the Phase 1 melt inclusions (Fig. 8). Water contents in melt inclusions from all Paricutin eruptive phases
range from 1·3 to 4·2 wt % (Luhr, 2001; Pioli et al., 2008),
and show no systematic variations during the eruption.
Several melt inclusions from the Phase 1 samples contain
measurable magmatic CO2 contents that range from 250
to 1000 ppm (Luhr, 2001; Pioli et al., 2008; Johnson et al.,
2009), whereas the later erupted material is essentially
devoid of CO2. This implies that olivine crystallization
took place over a range of pressures from 20 to 400 MPa
(51km to 13 km) in the Phase 1 magmas, whereas olivine
crystallization in the subsequent Phase 2 and 3 magmas
occurred at shallower levels (200 MPa; 6·6 km depth).
Therefore, Phase 1 magmas appear to have been erupted
from a deeper crustal level and possess distinctly different
major and trace element compositions relative to later
erupted material. New trace element data suggest either
that the Phase 1 and Phase 2 magma batches came from
compositionally different mantle sources or that the incompatible element differences result from assimilation in
NUMBER 11
NOVEMBER 2011
the lower crust of material that is distinct from the entrained xenoliths; this would require additional isotopic
data (Nd^Pb^O) to evaluate further.
Origin of Type I low-Si melt inclusions
(Eruptive Phases 2 and 3)
As previously documented, Type I inclusions have characteristically lower SiO2 concentrations relative to
whole-rock samples but otherwise have major and trace
element temporal variations that match those of the
whole-rock samples. The systematic offset between the
Type I inclusions and the whole-rock compositions probably reflects a combination of crystallization of the melt,
minor contamination, and crystal accumulation in the
magma after inclusion entrapment. Erlund et al. (2010)
observed that some Phase 2 melt inclusions (Luhr, 2001),
as well as early erupted bombs (Foshag & GonzalezReyna, 1956), similarly display systematically low SiO2
concentrations. Phase 2 melt inclusions from tephra samples also have Al2O3 concentrations higher than tephra
glass (Luhr, 2001), consistent with the observed
high-Al2O3 melt inclusions in olivine from the lava samples, but with similar CaO/Al2O3 ratios to the host
magma.
For the remainder of the discussion on melt evolution
and Type I inclusions we focus on trace element abundances in both whole-rock samples and melt inclusions.
Homogeneity of Eruptive Phase 2
samplesçany melt inclusion evidence
for assimilation?
Lavas and tephras erupted during Phase 2 (July 1943 to
1946) show limited major element variability, with MgO
of 4·7^5·3 wt % and SiO2 of 55·1^56·6 wt %. As noted by
Erlund et al. (2010), the Phase 2 tephra are remarkably
homogeneous in terms of whole-rock composition, mineral
compositions, and groundmass texture. Elemental ratios
such as K2O/TiO2 and Ba/Nb would be insensitive to crystal fractionation in the Paricutin magmas, but sensitive to
minor inputs of crustal material similar to the local basement and xenoliths. Compared with the preceding Phase
1 magmas, the Phase 2 magmas show greater variability
of K2O/TiO2 (1·0^1·3) and Ba/Nb (39^51), indicative of
minor amounts of crustal contamination (Fig. 11).
Variations in melt inclusion composition provide an
opportunity to evaluate the timing and amount of contamination relative to crystallization of the magma. For the
two Phase 2 lava samples analyzed, K2O/TiO2 ratios
(1·14 0·15 and 1·23 0·11) in Type I melt inclusions are
relatively constant and similar to whole-rock values (1·32
and 1·27 respectively). This is consistent with olivinehosted melt inclusions from the tephra with an average
K2O/TiO2 ratio of 1·18 0·09 compared with the
whole-rock tephra value of 1·27 0·02 (Luhr, 2001).
Similarly, Cl/K ratios are constant (0·08 0·015), despite a
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ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
range in SiO2 of 52^54 wt %. Only one melt inclusion
appears to record a more contaminated composition with
a K2O/TiO2 of 1·47 and Cl/K of 0·07. This inclusion also
has higher Ba/Nb (49·7) relative to the whole-rock and
other Phase 2 melt inclusions. As indices of contamination
(K2O/TiO2, Ba/Nb) have similar ranges in the wholerocks and melt inclusions, this would suggest that any significant contamination of the Phase 2 magma occurred
prior to olivine crystallization and entrapment of the Type
I melt inclusions. This is consistent with the generally low
crystallization pressures estimated for these magmas
based on H2O contents and a general lack of CO2 in
olivine-hosted melt inclusions (Luhr, 2001; Pioli et al.,
2008). Erlund et al. (2010) proposed a model by which the
Phase 2 magmas crystallize during ascent, with early
erupted magmas retaining some CO2 whereas later
erupted Phase 2 lavas record low crystallization pressures
(5100 MPa) and the establishment of a shallow magma
storage region.
Nature of the transition from Phase 2 to
Phase 3
The most distinctive compositional change in the Paricutin
lavas occurred at the end of 1946 and through 1947, with
a shift from basaltic andesite (56 wt % SiO2) to andesite
(60 wt % SiO2) compositions (Wilcox, 1954) that was
accompanied by significant increases in 87Sr/86Sr and
d18O, indicative of increased crustal assimilation
(McBirney et al., 1987; Cebria¤ et al., 2011). Several studies
have used the onset of this change to mark the transition
between the Phase 2 and Phase 3 eruptive stages (e.g.
Pioli et al., 2008; Erlund et al., 2010). However, our new
trace element analyses, combined with recent literature
data (Luhr, 2001; Erlund et al., 2010; Cebria¤ et al., 2011), indicate that, in many respects, there was a more fundamental compositional change almost a year earlier, sometime
in early 1946, which is used here to mark the onset of the
Phase 3 stage. This change is most clearly defined by a
shift in Zr/Nb from 15·4^18·9 in the Phase 2 magmas to
higher values (19·0^22·3) in the Phase 3 magmas that are
similar to those of the Phase 1 magmas (Fig. 11). Although
some of the scatter in Zr/Nb values is a result of slight
interlaboratory calibration differences, the observation of
a change in Zr/Nb at this time is a robust feature of the
data as it is apparent in three independent studies that
have analyses of both Phase 2 and Phase 3 samples (this
study; Luhr, 2001; Cebria¤ et al., 2011).
A striking feature of the lavas erupted between mid-1943
and mid-1947 that span this shift in Zr/Nb ratios is their
relatively constant MgO content (5·2^5·8 wt %), despite
small progressive increases in SiO2, K2O/TiO2, and
87
Sr/86Sr (Fig. 12). Compared with the earlier low Zr/Nb
samples, the subsequent high Zr/Nb samples have higher
SiO2 (55·4^56·6 wt % vs 56·7^57·8 wt %), K2O/TiO2
(1·1^1·3 vs 1·3^1·6), Ba/Nb (39^48 vs 52^65), Ba/La (21^22
vs 24^26) and Th/Nb (0·18^0·21 vs 0·22^0·26), and lower
Nb contents (7^8 ppm vs 8^10 ppm). Although the higher
SiO2, K2O/TiO2, Ba/Nb, Ba/La and Th/Nb are suggestive
of an increased input of crustal material, the constant
MgO contents, together with constant Ni (88^127 ppm),
Cr (153^227 ppm) and CaO/Al2O3 (0·39^0·41), make it difficult to explain such variations with a progressive assimilation and fractional crystallization model as proposed by
Wilcox (1954), McBirney et al. (1987) and Cebria¤ et al.
(2011). The local crustal lithologies (bulk xenoliths, xenolith glasses, basement outcrops) have average Zr/Nb in
the range 20^30, which does not give sufficient leverage to
account for the change in Zr/Nb in these lavas at similar
MgO values. Instead, these data suggest that the shift in
Zr/Nb values represents the involvement of a new magma
batch, compositionally distinct from the Phase 2 magma.
The minor variations in K2O/TiO2 and Ba/Nb, with limited variation in SiO2, seen in the later Phase 2 magmas
might be better explained by mixing with this new
magma batch rather than small extents of crustal
assimilation.
The exact timing of the input of this new high Zr/Nb
magma batch into the eruptive system at Paricutin is somewhat uncertain, as the Smithsonian collection has only a
few samples from the critical time period from 1945 to
1947. The recent study of Cebria¤ et al. (2011) improved
sample coverage in this interval, but their positions within
the eruptive chronology are known only to within
4^5 months. Samples Par-2 and Par-4 from Cebria¤ et al.
(2011) have high Zr/Nb and have inferred eruption dates
between October 1945 and February 1946, whereas sample
116289-8 (Table 1) has low Zr/Nb and was erupted on 18
September 1946. It was around this time that there was a
change in eruptive style from explosive Strombolian activity to effusive lava flows and Vulcanian explosions, accompanied by a decrease in the average mass eruption rate
(Pioli et al., 2008). Pioli et al. (2008) noted that this shift in
eruptive style preceded the rapid compositional change
from basaltic andesite to andesite compositions by several
months, and this means that it was potentially linked to
the appearance of the new high Zr/Nb magma batch.
Despite the rapid change in whole-rock compositions,
the melt inclusions do appear to preserve limited overlap
between the end of Phase 2 and the beginning of Phase 3,
implying minor mixing of these two phases at the onset of
Phase 3. Overlap in melt inclusion compositions is most
evident in Cl/K, Ba/Nb, and K2O/TiO2 ratios between
early Phase 3 and late Phase 2 (Figs 8 and 11). Ba/Nb
ratios in Phase 3 melt inclusions are as low as 45 compared
with a maximum Ba/Nb of 48 in Phase 2 melt inclusions.
Additionally, as discussed above, one inclusion in a late
Phase 2 lava has low Cl/K (0·7) and high K2O/TiO2
(1·5), comparable with the Phase 3 inclusions.
Importantly, this suggests the presence of Phase 3
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NUMBER 11
NOVEMBER 2011
Fig. 12. Whole-rock (a) SiO2, (b) Zr/Nb, (c) K2O/TiO2, and (d) Ba/Nb vs MgO wt %. The significant variations in indices of crustal contamination with little change in MgO concentration should be noted. Data sources: this study; Wilcox (1954); McBirney et al. (1987); Luhr (2001);
Verma & Hasenaka (2004); Erlund et al. (2010); Cebria¤ et al. (2011),
magmas within the shallow magma storage system during
the waning of Phase 2 activity. This model, incorporating
minor mixing between Phase 2 and Phase 3, is consistent
with the apparent overlap in whole-rock compositions
during this time interval.
Compositional variations of Phase 3
samples
Eruptive Phase 3 represents two-thirds of the eruption
duration (1946^1952) but only the last 25% of erupted
material (McBirney et al., 1987). Phase 3 lavas are characterized by a wide range in SiO2 (56·5 to 60·5 wt %) and
87
Sr/87Sr (Fig. 1). By comparison, Phase 2 lavas record the
bulk of the eruption (60%) but only record an 2 wt %
increase in SiO2. Phase 3b lava compositions are relatively
homogeneous, with most of the variation in Phase 3 occurring prior to August 1948 (Phase 3a). Phase 3 whole-rock
compositions are characterized by higher SiO2 concentrations and, based on indices of crustal contamination, generally appear more contaminated. Whole-rock Ba/Nb is
52^73 (vs 40^48 in Phase 2 lavas), and K2O/TiO2 is
1·2^2·2 (vs 1·2^1·3 in Phase 2), with both ratios increasing
with SiO2 and eruption date. The only exception is one of
the last erupted lavas (116289-19; 25 February 1952), which
has slightly lower SiO2 and lower Ba/Nb (69·3). Melt inclusions from this last sample are also hosted in more forsteritic olivine (average Fo80·1) compared with more fayalitic
olivine hosts (average Fo78·6) erupted earlier in Phase 3b,
suggesting that the last erupted lavas from Paricutin were
more primitive and less crustally contaminated than the
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ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
previously erupted material (Fig. 4). Variations in incompatible trace element ratios sensitive to crustal contamination (Ba/Nb, K2O/TiO2) in Phase 3 lavas and melt
inclusions can generally be modeled by mixing an early
Phase 3 lava (Par-2 of Cebria¤ et al., 2011) with a bulk xenolith composition similar to 116293-5 (Table 2). Using a
bulk xenolith composition, 25^30 wt % crustal addition is
required to explain the variations in trace elements
(Fig. 11). This assimilation model, coupled with minor fractional crystallization, can account for the majority of both
the major and trace element variability in Phase 3 melts.
The limited variations in Zr/Nb in Phase 3 magmas are
broadly consistent with assimilation of crust with similar
Zr/Nb, like the xenoliths. Scatter in Zr/Nb values (Figs 11
and 12) is most probably due to minor interlaboratory
biases, but minor mixing with Phase 2 melts could also
play a role.
Over the course of the eruption of Phase 3 lavas, Ba/Nb
in melt inclusions increases with decreasing host olivine
forsterite content, indicating that crystallization of more
evolved olivines is capturing more contaminated magma
(Fig. 13). K2O/TiO2 shows little systematic variation with
the forsterite content of olivine from Phase 3 lavas.
Similarly, Cl/K is negatively correlated with K2O/TiO2,
once again suggesting that decreasing Cl/K is recording
crustal assimilation of a high-K contaminant (Fig. 8b).
Despite geochemical evidence for contamination throughout the eruption of Phase 3 lavas, melt inclusions within
single lavas are similar to the whole-rock compositions
in terms of indices of contamination and display no
systematic variations with either host composition or melt
SiO2 within single samples. K2O/TiO2 in olivine-hosted
melt inclusions from the two Phase 3b lavas varies from
1·29 0·16 to 2·00 0·32, with corresponding wholerock ratios varying from 1·38 to 2·00, respectively.
Orthopyroxene-hosted melt inclusions from Phase 3 lavas
have K2O/TiO2 ratios equivalent to those of the
olivine-hosted melt inclusions, indicating a similar timing
of crystallization of olivine and orthopyroxene relative to
assimilation. Average Ba/Nb ratios in the melt inclusions
are generally offset to slightly lower values, but are within
error of the whole-rock values. These minor differences in
Ba/Nb between whole-rock and average melt inclusions
may be the result of minor late-stage contamination, following melt inclusion entrapment in olivine and prior to
formation of the Type II inclusions. However, because
Type II melt inclusions (excluding sample 116289-19) all
have Ba/Nb ratios equivalent to Type I inclusions and less
than the whole-rock values, the slight difference between
the melt inclusions and the whole-rock ratios is probably
an analytical artifact (reflecting slight calibration differences between SIMS and ICP-MS techniques). Therefore,
despite the evidence for significant crustal contamination
in the evolution of Phase 3 magmas, the melt inclusions
within single Phase 3 lavas do not preserve a record of progressive contamination and fractionation, implying that
any significant contamination must have occurred prior
to both Type I olivine and orthopyroxene crystallization.
However, in a given lava sample, variability in melt inclusion Ba/Nb ratios is outside analytical uncertainty
(Fig. 13); therefore, instead of a consistent and progressive
contamination signature, relatively enriched and variable
K2O/TiO2 and Ba/Nb (Fig. 11) ratios in Phase 3 melt inclusions may imply a relatively homogeneous ‘bulk’
magma experiencing variable contamination decoupled
from significant fractionation.
Magma chamber models and the timing of
inclusion entrapment and assimilation
Fig. 13. Olivine host forsterite content vs inclusion Ba/Nb ratio for
Type I melt inclusions.
Several physical models have been suggested for magma
chamber evolution and dynamics beneath Paricutin
Volcano as a means of explaining the observed compositional variations. Wilcox (1954) and McBirney et al.
(1987) both suggested that the eruptive history of the volcano was too short to have developed the observed compositional variations and therefore a stratified magma
chamber had to have been present for decades prior to
the eruption. In a stratified magma chamber model, density differences between compositionally distinct layers
halts convection between them, with contaminated
magma accumulating at the top of the chamber (Wilcox,
1954). However, convection within stratified layers could
still persist. McBirney et al. (1987) presented a similar
model in which the mafic intrusion melts the wall-rock
and then back-mixes with the crustal melt as it rises to
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collect under the roof. This model is also capable of generating the erupted volumes of zoned magma on a decadal
time scale depending on the shape and rate of cooling of
the intrusion. The model of a large zoned magma body as
the storage site for the Paricutin lavas, however, appears
to be fundamentally flawed in that it is based on the assumption that all of the erupted magmas are petrogenetically related by a simple progressive assimilation and
fractional crystallization process. This is not consistent
with the whole-rock trace element data, in particular incompatible trace element ratios, from both this study and
the literature that indicate that the eruption is composed
of at least two different magma batches. The distinction is
most easily observed in Zr/Nb ratios (Figs 11 and 12),
which show differences between Phase 1 and Phase 2
lavas and between Phase 2 and Phase 3 lavas that cannot
be explained by crustal assimilation. Additionally, a
pre-existing zoned magma chamber model does not explain the compositionally distinct Phase 1 lavas, which
record olivine crystallization at a significantly deeper
crustal level. The presence of short-lived, compositionally
distinct magma batches at the onset of the eruption is
hard to reconcile with a model dependent on a
pre-existing, stratified magma chamber, especially given
the lack of interaction between Phase 1 and Phase 2, as
indicated by both whole-rock data and melt inclusions.
Erlund et al. (2010) also argued against a pre-existing
zoned magma chamber based on the concept that some
of the melt inclusions from the early stages of the eruption
should have shown evidence for contamination, given
that most crystallization would occur in contaminated
thermal boundary layers.
An alternative model argues that there is no
pre-existing magma chamber and that basaltic andesite
magma is injected into the crust with subsequent rapid
crystallization and assimilation (Dungan, 2005;
Erlund et al., 2010; Cebria¤ et al., 2011). This model is based
on observations that crustal xenoliths can potentially be
assimilated rapidly (less than decadal time scales) and
that they would not survive in a hot basaltic andesite for
more than 10 years, as might be implied by the presence
of accumulated contaminated magma in a stratified
magma chamber (Wilcox, 1954; McBirney et al., 1987).
This is supported by the observation at Paricutin that
crustal xenoliths are found only in Phase 1 and 2 lavas,
and not in the high-SiO2, high-87Sr/86Sr Phase 3 lavas. In
the Dungan (2005) model, the eruption initially pulls material from the middle of the intrusion where crustal fragments have not yet been fully assimilated. Magma at the
edges of the intrusion continues assimilating crust and is
later mixed and erupted with the more primitive material.
In this scenario, the presence of an early, distinct magma
composition is easier to rationalize. However, using
mass-balance calculations McBirney et al. (1987)
NUMBER 11
NOVEMBER 2011
demonstrated that although an average xenolith composition is an appropriate contaminant for most of the trace
elements, none of the crustal xenoliths or basement samples are appropriate for explaining the Sr isotopic variability between Phases 2 and 3, which requires a
contaminant with a more radiogenic Sr isotope composition and greater Sr abundance. This requires a crustal
component not found in either the exposed basement or
the entrained crustal xenoliths: this is difficult to explain
if the magma is rapidly evolving by assimilation of its
local surroundings as suggested by Dungan (2005).
Additionally, assuming that the melt inclusions provide a
representative sampling of the evolving magma, there is
no melt inclusion evidence for this rapid, progressive
magma evolution. In the Dungan (2005) model, one of
the key requirements is that the same basaltic andesite is
essentially present throughout the eruption, with erupted
magma compositions varying as a result of rapid, late assimilation of crustal xenoliths entrained in the magma. If
this model is correct, we might expect to find melt inclusion compositions in the Phase 3 lavas that record a mixture of both contaminated and uncontaminated (Phase 1
or 2) magma batches. Although Type I melt inclusions in
the first sample in Phase 3 do record a wide range of Ba/
Nb ratios the inclusions predominantly tend to record discrete compositions more characteristic of crystallization
within an already contaminated magma body. Erlund
et al. (2010) presented a model similar to that of Dungan
(2005), in which during high mass flux eruption rates
deeper magmas (as indicated by measurable CO2 concentrations) fed the eruption and caused sill formation.
Magma stored at shallow depths in dikes and sills would
then erupt subsequently as mass flux rates from depth
decreased. This model is attractive in that it provides an
explanation for the initial compositionally distinct Phase
1 magmas.
Relative timing of AFC processes and
melt inclusion entrapment
Johnson et al. (2008) suggested for Volcan Jorullo (Mexico)
that whereas melt inclusions record crystallization of olivine during ascent and degassing of the magma, the compositional trends defined by lavas are driven by a deeper
fractionation that is not recorded in the erupted phenocrysts and inclusions. Essentially the crystal^melt inclusion
record is missing the deeper processes. A similar model
may be applicable to Paricutin and would be consistent
with the generally low crystallization pressures of Phase 2
and 3 lavas estimated from water concentrations (and no
detectable CO2) in melt inclusions (Luhr, 2001). This
would provide a mechanism to explain whyType I melt inclusions have a restricted compositional range similar to
the whole-rock composition for single samples rather than
a range of compositions from primitive to more evolved as
2214
ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
would be expected as a consequence of a progressive AFC
process.
An alternative interpretation of the inclusion data is that
assimilation and fractional crystallization were decoupled.
One of the fundamental underpinnings of the coupled
assimilation^fractional crystallization (AFC) model (e.g.
DePaolo, 1981) is that the latent heat of crystallization
during fractional crystallization provides the heat source
for crustal assimilation, such that magmas evolve through
progressive and concurrent assimilation and fractional
crystallization as observed in olivine-hosted melt inclusions
by Kent et al. (2002) in Yemen flood basalts. If this model
is correct, melt inclusions and phenocrysts in crustally contaminated magmas should record a progression from a
primitive uncontaminated magma toward a contaminated
more evolved magma. However, this progressive contamination is not recorded by melt inclusions at Paricutin.
Instead, ParicutinType I melt inclusions record a relatively
restricted compositional range in each of the sampled
lavas (Fig. 4; Table 5). This is most evident when comparing
inclusions between eruptive phases. In a progressive AFC
model, inclusions in Phase 3 lavas might be expected to
record a range of compositions, from more primitive
(Phase 2) to more evolved and contaminated. However,
Type I inclusions (the most primitive) in Phase 3 lavas
appear to be more contaminated relative to Phase 2 inclusions and record a restricted compositional range with
only minor overlap with Phase 2 compositions at the
onset of Phase 3 (Figs 8 and 11). This may suggest that contamination in the Paricutin magmatic system is decoupled
from crystallization, or at least from the crystallization recorded by melt inclusions in the lavas. A similar interpretation has previously been suggested by Grove et al. (1988)
to explain the generation of an andesite lava at Medicine
Lake, California. In this example, the arguments for
decoupling of crystallization and assimilation are based
on petrological evidence rather than a direct record of
magma compositions, and the observation that in silicate
melts, diffusion of heat is substantially faster than diffusion
of chemical constituents.
One mechanism to explain the apparent decoupling of
assimilation and fractional crystallization in the melt inclusion record at Paricutin is that the thermal conditions
required for inclusion entrapment were not favorable to
coupled crustal assimilation and fractional crystallization.
In this case the melt inclusions represent a biased record
of magma evolution with much of the crystallization and
assimilation taking place within a main magma chamber
or body essentially missed by the inclusion record.
Petrographic investigations and experimental studies of
melt inclusion formation have identified several ways of
forming melt inclusions, with isothermal crystallization
following rapid cooling (e.g. Roedder, 1979, 1984; Kohut &
Nielsen, 2004) and entrapment in crystal defects or
dislocations during slower cooling (e.g. Faure & Schiano,
2005), which is probably the most realistic for basaltic compositions (Kent, 2008). Similarly, Goldstein & Luth (2006)
found that inclusions formed at a range of cooling rates
(7^2508C h1) but not at very low cooling rates (18C h1;
Kent, 2008). The potential need for undercooling to form
melt inclusions may preclude significant assimilation and
may be related to magma chamber processes, particularly
where in the magma chamber the thermal conditions are
favorable for inclusion formation and entrapment relative
to wallrock assimilation. If this is the case, melt inclusions
in short-lived, small-volume magmatic systems might provide a biased sampling of magma chamber processes and
melt evolution.
Multiple magma batches and magma
mixing
The preferred explanation for the overall evolution of the
Paricutin magma system is the presence of several compositionally distinct magma batches at shallow levels beneath
the erupting volcano. Abrupt temporal changes in ratios
of highly incompatible trace element ratios such as Zr/Nb
that are difficult to explain using a progressive assimilation
and fractional crystallization model mark the arrival of
distict magma batches. The Phase 1 and Phase 2 magmas
have similar 87Sr/86Sr but are compositionally different in
terms of Zr/Y ratios (6 vs 7^8) and LREE^HREE fractionation (La/YbN: 5·3^5·7 vs 6·6^7·2), and these features
can be explained by small differences in the degree of melting to produce distinct magma batches. The transition
from Phase 2 to Phase 3 records the shift and potential
minor mixing between two independently evolving but
interconnected magma bodies. This model was also
favored by Luhr and Housh based on mineral phase equilibria of the Phase 2 and Phase 3 lavas (T. Housh, personal
communication, 2007). The Phase 3 magmas represent
only a small proportion of the total erupted volume, and
they record a wide compositional range that is consistent
with crustal assimilation and fractional crystallization.
The relationship, if any, of the Phase 3 magmas to the two
preceding magma batches is uncertain. Even the least contaminated Phase 3 sample has higher 87Sr/86Sr and K2O/
TiO2 than any of the earlier Phase 1 and 2 magmas, indicative of some crustal influence. The difference in Zr/Nb
at similar MgO makes it difficult to relate the Phase 3
and Phase 2 magmas. On some trace element plots (e.g.
Fig 11), the composition of Phase 3 lavas could potentially
be explained by assimilation of an average xenolith composition by Phase 1 lavas, although this is not the case
with other trace element ratios (e.g. Fig. 14). Work is currently under way to assess the relationships between these
different magma batches using high-precision Pb isotope
analyses.
The evolution of Paricutin volcano appears to have
been
controlled
by
the
eruption
of
three
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JOURNAL OF PETROLOGY
VOLUME 52
NUMBER 11
NOVEMBER 2011
Fig. 14. (a) Nb/Yb vs Th/Yb diagram, after Pearce & Peate (1995), demonstrating the variability of magmatic compositions at Paricutin compared with those of the MGVF. Analyses of crustal material have been placed into one of four categories with the average composition for
each category plotted: (1) xenolith glass; (2) basement; (3) high 87Sr/86Sr xenoliths; (4) low 87Sr/86Sr xenoliths. Compiled MGVF data sources:
Luhr & Carmichael (1985); Hasenaka & Carmichael (1987); Luhr (2001); Chesley et al. (2002); Righter et al. (2002); Verma & Hasenaka (2004);
Johnson et al. (2009); Cebria¤ et al. (2011).
compositionally distinctive and independently evolving
magma batches. The existence of compositionally distinct
magma batches that are erupted at different times during
a monogenetic eruption is not unique to Paricutin.
Numerous studies of small-volume basaltic systems have
shown the eruption of compositionally distinct magma
batches over relatively short time periods (e.g. Camp
et al., 1987; Reagan & Gill, 1989; Cervantes & Wallace,
2003; Strong & Wolff, 2003). Reagan & Gill (1989)
observed Nb-enriched and Nb-depleted magmas from the
same eruptive episode at Turrialba Volcano, Costa Rica,
and argued for the tapping and flux melting of compositionally distinct mantle source domains. Similarly, Strong
& Wolff (2003) noted the contemporaneous eruption of
calc-alkaline and ocean island basalt (OIB)-like magmas
in southern Cascades monogenetic centers. They argued
that the different melts required distinct magma sources,
and that either the melts travel through the crust without
storage within a magma chamber or that the magma
chambers are inefficient at mixing these compositionally
distinct magma batches.
There is good evidence for the eruption of compositionally diverse magmas in close proximity in the
Michoacan^Guanajuato volcanic field (MGVF) where
Paricutin is located (Luhr & Carmichael, 1985; Verma &
Hasenaka, 2004; Johnson et al., 2009). Luhr & Carmichael
(1985) showed that magmas of considerably different composition were erupted only 3 km apart, albeit at different
times, at Volcan Jorullo (K2O/TiO2 1·0; La/Sm 3^4) and
Cerro la Pilita (K2O/TiO2 2·1; La/Sm 10^12). Johnson
et al. (2009) noted that there is no systematic across-arc
variation in magma composition within the MGVF.
Figure 14 shows how the compositions of the Phase 1, 2
and 3 magmas at Paricutin compare with compositional
diversity found within the MGVF, using a plot of Nb/Yb
vs Th/Yb (after Pearce & Peate, 1995).
CONC LUSIONS
The two distinct melt inclusion populations preserved in
single lava samples provide new insights into the history of
magma evolution and the relative timing of crystallization
and assimilation at Paricutin. The Type II (high-SiO2)
melt inclusions record a very late-stage crystallization
event with little assimilation and are compositionally similar to tephra matrix glasses reported by Luhr (2001) and
2216
ROWE et al.
MAGMATIC PLUMBING SYSTEM, PARICUTIN
Erlund et al. (2010). Because this population has low volatile contents and is not observed in contemporaneous
tephra samples, we suggest that these high-SiO2, low-S inclusions record very shallow crystallization after segregation of gases from denser magma near the base of the
edifice.
The low-SiO2, high-S population (Type I) of inclusions
records a relatively restricted compositional range, more
similar to the whole-rock compositions. If the melt inclusion compositions record a representative sampling of
magma chamber processes, none of the previous models
for magma chamber development and melt evolution at
Paricutin are appropriate. The similarity of Type I melt inclusions to single whole-rock compositions may instead
imply that crystallization, or at least the crystallization recorded by the erupted crystal cargo, must have occurred
after significant crustal contamination. The potential
decoupling of assimilation and fractional crystallization
may also be a function of the mechanism for inclusion entrapment in that inclusions may be forming only at specific
intervals in the magmatic evolution when thermal conditions (i.e. degree of undercooling) are optimal for inclusion entrapment.
In conjunction with the melt inclusion data, new
whole-rock trace element analyses indicate that the abrupt
compositional variations in incompatible trace element
ratios observed in erupted lavas at Paricutin are probably
the result of multiple, independently evolving, small
magma batches rather than the progressive assimilation
and fractional crystallization of a single batch of basaltic
andesite magma. Compositionally diverse magmas erupting in close proximity to one another are observed elsewhere locally in the Michoaca¤n^Guanajuato volcanic
field and have been attributed to mantle heterogeneity.
Therefore, although fractional crystallization and crustal
assimilation may be important processes within single
eruptive phases, the complexity and timing of compositional shifts in the erupted magma at Paricutin are instead
a function of the co-evolution of multiple, compositionally distinct magma batches derived by variable melt
generation processes within a heterogeneous mantle
source region, with additional crustal assimilation
effects superimposed prior to their input to shallow crustal
levels.
AC K N O W L E D G E M E N T S
The authors would like to thank Frank Tepley (Oregon
State University EMPA laboratory) and Rick Hervig and
Linda Williams (Arizona State University SIMS facility)
for their assistance. We thank Paul Wallace, Kate
Saunders, and an anonymous reviewer for feedback and
comments. Whole-rock samples and thin sections for this
study were provided on loan from the Smithsonian
National Museum of Natural History.
FU N DI NG
Funding for this project came from National Science
Foundation, Division of Earth Sciences, grant 0609652.
The SIMS facility at ASU is partly supported by a grant
from the Instrumentation and Facilities Program,
Division of Earth Sciences, National Science Foundation.
S U P P L E M E N TA RY DATA
Supplementary data for this paper are available at Journal
of Petrology online.
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APPENDIX
Fig. A1. (a) Location map for Paricutin volcano within the
Michoacan^Guanajuato volcanic field (MGVF) of central Mexico;
(b) map of the Paricutin region showing the local distribution of
cinder cones; modified from McBirney et al. (1987) and Erlund et al.
(2010). The plate configuration is also indicated in (a). P, Paricutin volcano; RP, Rivera Plate.
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