An Investigation into the Nature of the Magmatic Plumbing System at ...

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. ROWE 1,2 *, DAVID W. PEATE 1 AND 

INGRID UKSTINS PEATE 1 

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 

*Corresponding author.Telephone: (509) 335-6770. 

E-mail: mcrowe@wsu.edu 

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. 

KEY WORDS: trace element; magma chamber; melt inclusion; crustal 

contamination; crystallization 

INTRODUCTION 

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 % SiO 2: Fig. 1a; e.g. Wilcox, 

1954), accompanied by an increase in 87 Sr/ 86 Sr (Fig. 1b) 

and d 18 O 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). 

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JOURNAL OF PETROLOGY VOLUME 52 NUMBER 11 NOVEMBER 2011 

Fig. 1. (a) Eruption date vs SiO 2 (whole-rock lava and tephra data); (b) inset plot of SiO 2 vs 87 Sr/ 86 Sr. 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 

2188 

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 



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. 

ERUPTIVE HISTORY OF 

PARICUTIN VOLCANO 

Paricutin lies within the Michoaca¤n^Guanajuato volcanic 

field (MGVF), which contains over 1000 small eruptive 

centers (Appendix Fig. A1) over an 40 000 km 2 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 km 3 . 

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 K 2O 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 % SiO 2; 87 Sr/ 86 Sr 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 % SiO 2; 87 Sr/ 86 Sr 0·7042; McBirney et al., 1987). 

Crustal material, either as xenoliths or xenocrysts, is very 

rare in the Phase 3 eruptive material. 

2189 

SAMPLE DETAILS AND 

ANALYTICAL METHODS 

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 inTable 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). 




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 

2190 

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 CO 2^H 2 gas mixture. Total time of heating above 

10008Cwaskeptto 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 


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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 %) 

SiO 2 54·59 55·39 55·71 55·79 56·13 57·05 58·39 59·09 

TiO 2 0·99 0·94 1·01 0·90 1·02 0·89 0·86 0·78 

Al 2O 3 17·83 17·64 17·24 17·48 17·34 17·27 17·78 17·55 

Fe 2O 3 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 

Na 2O 3·95 3·98 3·99 3·81 3·79 3·71 3·86 3·92 

K 2O 0·91 1·15 1·18 1·19 1·30 1·23 1·30 1·50 

P 2O 5 0·27 0·35 0·33 0·30 0·36 0·29 0·30 0·08 

H 2O þ 


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 0·29 

Ba 311 372 388 400 398 416 472 506 

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 7·45 

Rb 10·4 15·3 16·6 17·0 17·3 19·0 20·6 22·4 

Cs 0·35 0·43 0·42 0·44 0·45 0·53 0·60 0·58 

Sr 603 588 594 582 576 565 562 547 

Sc 18·1 16·6 17·7 17·2 18·0 17·3 14·5 14·2 

Zr 109 146 157 146 152 135 145 155 

2191 

(continued) 



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 %) 

SiO 2 59·77 60·24 60·07 71·93 70·88 72·61 71·00 75·95 72·47 

TiO 2 0·83 0·80 0·81 0·23 0·36 0·17 0·18 0·04 0·20 

Al 2O 3 17·29 17·3 17·28 14·98 14·27 14·98 14·83 13·51 14·51 

Fe 2O 3 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 

Na 2O 3·74 4·01 4·00 4·67 4·18 4·75 4·13 3·90 4·33 

K 2O 1·67 1·66 1·67 2·15 3·64 1·63 2·38 4·74 2·91 

P 2O 5 0·12 0·04 0·03 0·19 0·11 0·39 0·47 0·13 0·14 

H 2O þ 


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 

Total 100 99·95 99·94 99·9 99·49 100·16 99·08 99·95 99·72 

Ni (ppm) 57 44 63 11 18 15 

Cr (ppm) 67·8 66·5 76 4·7 18 15 37 10 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 0·27 0·28 0·26 0·14 0·28 0·14 0·13 0·31 0·20 

Ba 540 552 554 327 498 339 676 124 393 

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 24·7 26·0 27·2 44·8157·3 34·0 62·2 149·9 89·7 

Cs 0·64 0·67 0·72 1·17 2·30 1·14 1·23 2·82 1·73 

Sr 531 528 534 395 50 441 458 52 279 

Sc 14·5 15·0 14·4 2·6 8·5 2·9 2·8 1·9 3·72 

Zr 161 161 158 106 153 107 102 45 102 

2192 




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.) 

2193 

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 inTable 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 ( 88 Sr, 89 Y, 90 Zr, 93 Nb, 138 Ba, 

139 La, 140 Ce, 144 Nd, 147 Sm, 151 Eu, 158 Gd, 162 Dy, 174 Yb) were 

analyzed on a Cameca 3f ion microprobe, together with 

30 Si and 42 Ca for calibration and to allow correction for 

potential olivine overlap. For trace element analysis we 

used a 16 O 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 88 Sr to 139 La and 140 Ce to 174 Yb) with 30 Si 

measured before and after each block for normalization. 

Count times were 10 s ( 30 Si), 20 s ( 42 Ca, 88 Sr), 30 s (masses 

from 89 Yto 144 Nd, plus 158 Gd and 162 Dy) or 40 s ( 147 Sm, 

151 Eu, 166 Er, 174 Yb). Measured ratios (M þ / 30 Si) were corrected 

for interfering oxides using rare earth element 

(REE) oxide production values (MO þ /M þ ) from Zinner 

& Crozaz (1986) and a 135 Ba 16 O/ 135 Ba oxide production 

ratio of 0·054 (R. Hervig, personal communication). For 

small inclusions, CaO concentrations were calculated 

from 42 Ca/ 30 Si 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 144 Nd 

and 6^9% for masses from 147 Sm to 174 Yb. Accuracy, 

relative to preferred values for BCR-2G (GEOREM: 

http://georem.mpch-mainz.gwdg.de), is better than precision 

for all elements except 158 Gd (þ11%) and 174 Yb 

( 13%). Additional details for accuracy and precision for 

basaltic trace element analysis by SIMS are presented in 

Table 4. 




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 15·4 

F 0·01 – – 0·02 – – 0·01 – – – – – – – – 

Total 98·90 99·34 – – 97·40 – – 99·82 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, 47 Ti and 232 Th were also 

counted. We used a 16 O 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 144 Nd were counted for 1s 

( 93 Nb, 139 La, 144 Nd counted for 4, 2, and 4 s, respectively), 

and ions with a mass greater than 147 Sm were counted for 

5 s for each measurement cycle (25 cycles per analysis). 

Measured M þ / 30 Si 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%). 

2194 

RESULTS 

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 (K D) of 0·30 is 


http://petrology.oxfordjournals.org/ 

by guest on April 3, 2013


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 

Ba 131·0 3·9 – 686·7 1·9 0·5 476·2 10·1 10·9 

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 10·8436·1 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 13·0466·3 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 K D value 0·34 0·02 based on olivine core compositions 

and bulk tephra Mg-number using the method described 

by Toplis (2005). However, a K D 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 KD Fe^Mg of 0·32. Orthopyroxene^ 

melt K D Fe^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 K D Fe^Mg , as this process should produce a negative 

correlation between these two parameters in a suite of 

2195 

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 K d values as 

described above and Fe 3þ /Fe total 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 SiO 2) vary by less than 5%, 

whereas MgO and FeO vary by up to 30%. Only the 




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. 

2196 

Corrected major element compositions of 

melt inclusions 

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 SiO 2 

(Type I) and one with higher SiO 2 (Type II), with a typical 

average gap of 4^8 wt % SiO 2 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 

SiO 2 and Al 2O 3. Type I inclusions have SiO 2 concentrations 

equivalent to whole-rock SiO 2 concentrations in 

Phase 1 lavas, but SiO 2 is lower relative to whole-rock compositions 

by up to 4 wt % in eruptive Phases 2 and 3 

(Fig. 4). Al 2O 3 and CaO concentrations are proportionally 

enriched in all inclusions relative to whole-rock abundances, 

from 0·3 to2·1 wt % and from 0·3 to0·7 wt%,respectively 

(Fig. 6). Generally, TiO 2 and P 2O 5 contents in 

the melt inclusions decrease over the course of the eruption, 

similar to the whole-rock trends (Fig. 6). Al 2O 3 concentrations 

do not vary systematically and remain 

relatively constant throughout Phase 3. K 2O and SiO 2 are 

the only major elements to increase in theType I inclusions 

over the course of the eruption, with average K 2O 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. SiO 2 concentrations in Type I inclusions 

are lowest in Phase 2 lavas with average SiO 2 as 

low as 52·5 wt %. From Phase 2 to Phase 3, SiO 2 increases 




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 SiO 2,TiO 2,K 2O, and P 2O 5, relative to both 

Table 5: Average corrected melt inclusion major and trace element compositions 

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 TiO 2 vs 1·2 wt % average TiO 2 in 

Phase 2). The average concentrations of both CaO and 

Al 2O 3 increase through Phase 1 and 2 and then decrease 

significantly in Phase 3 lavas (Table 5). However, at 

all times average CaO and Al 2O 3 contents in Type II 

inclusions are lower than those of the whole-rock and 

Sample: 116293-7 116293-7 116295-23 116295-23 116289-8 

Comment: Type I Type II Type I Type II Type I 

Host: Olivine Olivine Olivine Olivine Olivine 

Av. SD Av. SD Av. SD Av. SD Av. 

Corrected inclusion composition (wt %) 

SiO2 54·22 1·09 61·93 0·88 52·71 2·95 56·52 1·24 52·51 

TiO2 1·24 0·11 2·17 0·49 1·06 0·38 1·18 0·29 1·11 

Al2O3 20·29 0·73 15·47 1·61 17·79 1·66 16·82 2·20 18·41 

FeO 4·60 0·69 4·39 0·97 6·91 1·30 5·97 1·34 6·65 

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 

Total 100·00 0·00 100·00 0·00 100·00 0·00 100·00 0·00 100·00 

Host Mg-no. 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 0·06 

Dilution 1·04 1·04 0·99 1·03 1·01 

Trace elements (ppm) 

Sr 668 51 440 153 580 0·4 595 171 507 

Y 18·5 1·2 24·9 12·3 18·9 0·4 17·5 8·7 24·1 

Zr 121 7 186 89 146 1·1 135 76 165 

Nb 6·3 0·5 9·8 3·6 8·9 0·6 8·6 4·4 11·2 

Ba 351 41 465 30 395 5 405 91 555 

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 

Nd 17·7 1·9 23·7 10·6 20·8 1·6 19·4 9·1 25·8 

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 

2197 

(continued) 





Sample: 116289-8 116289-9 116289-9 116289-12 116289-12 116289-15 

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. SD Av. SD 


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 – 0·99 – 1·03 – 


Sr 36 585 44 389 37 593 30 287 26 542 35 

Y 2·7 20·1 4·2 28·6 3·6 20·9 3·5 29·0 2·9 19·0 0·7 

Zr 33 154 26 214 36 168 33 253 16 167 7·0 

Nb 2·0 8·7 1·6 14·1 2·2 8·9 1·8 13·4 0·8 8·5 0·5 

Ba 35 454 63 752 52 515 75 812 83 565 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 

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 

2198 

(continued) 

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 SiO 2 





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 


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 – 

Total 100·00 0·00 100·00 0·00 100·00 – 100·00 0·00 100·00 0·00 100·00 0·00 100·00 – 

Host Mg-no. 78·17 0·26 80·77 1·70 81·57 – 80·09 0·27 79·41 0·84 79·11 3·27 77·91 – 

X (host) wt % 1·70 0·14 2·24 1·70 3·40 – 1·52 1·01 1·85 3·52 2·40 – 

Dilution 1·00 1·04 0·99 1·04 1·01 1·02 1·03 – 


Sr – – – – – – 523 32 396 137 – – – – 

Y – – – – – – 17·9 0·1 23·6 4·9 – – – – 

Zr – – – – – – 163 6 209 73 – – – – 

Nb – – – – – – 9·0 0·3 13·5 5·8 – – – – 

Ba – – – – – – 589 26 1093 164 – – – – 

La – – – – – – 19·2 1·3 28·6 5·3 – – – – 

Ce – – – – – – 41·2 2·8 58·7 10·7 – – – – 

Nd – – – – – – 21·6 1·4 28·2 6·2 – – – – 

Sm – – – – – – 4·2 0·5 6·1 1·0 – – – – 

Eu – – – – – – 1·0 0·1 1·5 0·6 – – – – 

Gd – – – – – – 3·5 0·5 5·7 3·0 – – – – 

Dy – – – – – – 3·3 0·7 4·9 1·5 – – – – 

Yb – – – – – – 1·7 0·2 2·0 0·4 – – – – 

*Compositions corrected using the spreadsheet and procedure described by Danyushevksy et al. (2000). 

contents, although the low-SiO 2 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 SiO 2 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 

2199 

Type I opx-hosted and olivine-hosted inclusions (Table 5). 

Higher FeO concentrations inType 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 




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). 

2200 




Fig. 5. (a) SiO 2 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 

2201 

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)andTh,withnegativeanomaliesat 

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/Sm N ¼ 2·63 0·23), elevated 

middle REE (MREE) to heavy REE (HREE) levels 




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/Yb N 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 

SiO 2, 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 

2202 

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 

Fo 83^76 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 Fo 85^87 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 




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 betweenType 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 




decrease over the course of the eruption from 0·12 to 0·06 

in Type I inclusions (Fig. 8). The negative correlation between 

K 2O/TiO 2 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 

K 2O concentrations as a result of the high K 2O 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 SiO 2 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 Fo 80·0 to Fo 83·5; Luhr(2001) 

reported a similar range in olivine compositions adjacent 

to melt inclusions in tephra samples from 1943 to 1948 

(Fo 80·0 to Fo 84·4). Average olivine compositions hosting 

Type II inclusions extend to more evolved compositions, 

ranging from Fo 82·0 to Fo 77·5. Olivine compositions hosting 

Type I and Type II inclusion populations have similar temporal 

variations. Average Phase 1 olivines (Type I ¼ Fo 80·8, 

Type II ¼Fo77·5) are more evolved than later Phase 2 olivines 

(Type I ¼ Fo 82·9,TypeII¼Fo 81·6). 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 Fo 88·4. 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 Fo 80·9, Type II¼Fo 79·4; 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 Fe T) 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 

2204 

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 % SiO 2), 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 Al 2O 3 and Sr and lower K, Rb, Y, HREE, Th and 

87 Sr/ 86 Sr 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 (SiO 2 

71^79%), with only a few dacite analyses (SiO 2 65^68%). 

In each xenolith, the glass compositions have higher K 2O, 

lower Na2O and much lower Th contents than the bulk 

composition, but otherwise exhibit similar compositional 

features. For example, xenolith 116289-23 has higher K 2O 

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. 

K 2O/TiO 2 ratios are similar in the glasses and bulk xenoliths, 

ranging from 5 to 15, except where TiO 2 contents 

are 50·1 wt %, when higher ratios (50^190) are observed. 




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, 

2205 

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 

SiO 2 concentrations, typically record a greater compositional 

range than the Type I inclusions within a particular 

whole-rock sample. They are displaced to higher TiO 2, 

P 2O 5 and K 2O values and lower Al 2O 3 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; 




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 64·13 

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 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 0·015 

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 98·54 


Ti 15·3 18·4 14·1 17·2343 92·8 253 223 1642 47·9 23·11848 9497 

Sr 11·2 11·3 9·1 99·1129 1016 210 135 291 280 467 226 445 

Y 6·3 2·9 1·8 5·7 5·1 13·2 5·2 4·8 16·2 16·6 2·8 10·3 81·0 

Zr 2·4 0·1 1·1 0·1 31·2 1·7562 30·3 24·8370 5·2 93·3 1003 

Nb 0·3 0·1 0·1 0·1 2·1 0·6 1·7 1·9 7·4 0·3 0·1 10·7 75·2 

Ba 43·0 53·3 31·7250 420 867 691 408 428 576 670 510 676 

La 1·3 1·1 1·5 1·4 6·8151 6·6 5·4 9·1 4·2 3·3 10·9 46·2 

Ce 1·4 1·7 2·3 1·9 12·2227 11·9 8·9 16·4 7·7 5·1 17·7 101 

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 (Fo 82·0^77·5) 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 K 2O and TiO 2) from the main temporal 

trend through the bulk lava and tephra samples (e.g. 

increasing K 2O and decreasing TiO 2). 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 K 2O/TiO 2 of each 

2206 

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) 




are lower (Fig. 4), consistent with the lower CaO and 

Al 2O 3 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 K 2O/TiO 2 (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 (K 2O/TiO 2 0·8^2·1). Therefore, contaminated 

magmas will have higher K 2O/TiO 2 values, and yet 

K 2O/TiO 2 does not significantly increase with SiO 2 in 

Type II inclusions within a single sample (Fig. 10). 

Although there is no direct correlation with SiO 2, the 

range in K 2O/TiO 2 values at constant SiO 2 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). 

2207 




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 betweenType 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 

2208 

2 magmas, highlighted by a clear gap in K 2O 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 SiO 2 

compared with the Phase 2 lavas, whereas Phase 1 melt inclusions 

on average have higher SiO 2 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 SiO 2 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 




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). 

2209 




indistinguishable 87 Sr/ 86 Sr ratios of Phase 1 and Phase 2 

lavas despite the elevated 87 Sr/ 86 Sr 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, K 2O/TiO 2, Sr/Nd) in olivine-hosted melt inclusions 

(Type I) show no systematic evidence of progressive 

contamination with melt evolution. K 2O/TiO 2 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 K 2O 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, SiO 2 or K 2O/TiO 2 

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 K 2O 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 to4·2 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 CO 2 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; 6·6kmdepth). 

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 

2210 

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 SiO 2 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 & Gonzalez- 

Reyna, 1956), similarly display systematically low SiO2 

concentrations. Phase 2 melt inclusions from tephra samples 

also have Al 2O 3 concentrations higher than tephra 

glass (Luhr, 2001), consistent with the observed 

high-Al 2O 3 melt inclusions in olivine from the lava samples, 

but with similar CaO/Al 2O 3 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 SiO 2 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 K 2O/TiO 2 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 K 2O/TiO 2 (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, K 2O/TiO 2 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 




range in SiO 2 of 52^54 wt %. Only one melt inclusion 

appears to record a more contaminated composition with 

aK 2O/TiO 2 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 

(K 2O/TiO 2, 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 H 2O contents and a general lack of CO 2 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 CO 2 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 % SiO 2) compositions (Wilcox, 1954) that was 

accompanied by significant increases in 87 Sr/ 86 Sr and 

d 18 O, 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 SiO 2, K 2O/TiO 2, and 

87 Sr/ 86 Sr (Fig. 12). Compared with the earlier low Zr/Nb 

samples, the subsequent high Zr/Nb samples have higher 

SiO 2 (55·4^56·6 wt % vs 56·7^57·8 wt %), K 2O/TiO 2 

(1·1^1·3 vs1·3^1·6), Ba/Nb (39^48 vs 52^65), Ba/La (21^22 

2211 

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 

SiO 2,K 2O/TiO 2, 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/Al 2O 3 (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 K 2O/TiO 2 and Ba/Nb, with limited 

variation in SiO 2, 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 K 2O/TiO 2 

( 1·5), comparable with the Phase 3 inclusions. 

Importantly, this suggests the presence of Phase 3 




Fig. 12. Whole-rock (a) SiO 2,(b)Zr/Nb,(c)K 2O/TiO 2, 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 SiO 2 ( 56·5 to60·5 wt %) and 

87 Sr/ 87 Sr (Fig. 1). By comparison, Phase 2 lavas record the 

bulk of the eruption ( 60%) but only record an 2wt% 

increase in SiO 2. Phase 3b lava compositions are relatively 

2212 

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 SiO 2 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 K 2O/TiO 2 is 

1·2^2·2 (vs 1·2^1·3 in Phase 2), with both ratios increasing 

with SiO 2 and eruption date. The only exception is one of 

the last erupted lavas (116289-19; 25 February 1952), which 

has slightly lower SiO 2 and lower Ba/Nb (69·3). Melt inclusions 

from this last sample are also hosted in more forsteritic 

olivine (average Fo 80·1) compared with more fayalitic 

olivine hosts (average Fo 78·6) erupted earlier in Phase 3b, 

suggesting that the last erupted lavas from Paricutin were 

more primitive and less crustally contaminated than the 




previously erupted material (Fig. 4). Variations in incompatible 

trace element ratios sensitive to crustal contamination 

(Ba/Nb, K 2O/TiO 2) 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). K 2O/TiO 2 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 

Fig. 13. Olivine host forsterite content vs inclusion Ba/Nb ratio for 

Type I melt inclusions. 

2213 

systematic variations with either host composition or melt 

SiO 2 within single samples. K 2O/TiO 2 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 K 2O/TiO 2 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 

K 2O/TiO 2 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 

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 




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-SiO 2,high- 87 Sr/ 86 Sr 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) 

2214 

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 CO 2) 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 




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 

2215 

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 h 1 ) but not at very low cooling rates (18C h 1 ; 

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 87 Sr/ 86 Sr but are compositionally different in 

terms of Zr/Y ratios ( 6 vs 7^8) and LREE^HREE fractionation 

(La/Yb N:5·3^5·7 vs6·6^7·2), 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 87 Sr/ 86 Sr and K 2O/ 

TiO 2 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 




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 87 Sr/ 86 Sr xenoliths; (4) low 87 Sr/ 86 Sr 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 

2216 

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 (K 2O/TiO 2 

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). 

CONCLUSIONS 

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-SiO 2) 

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 




Erlund et al. (2010). Because this population has low volatile 

contents and is not observed in contemporaneous 

tephra samples, we suggest that these high-SiO 2, low-S inclusions 

record very shallow crystallization after segregation 

of gases from denser magma near the base of the 

edifice. 

The low-SiO 2, 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. 

ACKNOWLEDGEMENTS 

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. 

2217 

FUNDING 

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. 

SUPPLEMENTARY 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. 

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RP, Rivera Plate. 

2220

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