JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 3
PAGES 421±456
2003
Capture of Helium and Other Volatiles
during the Growth of Olivine Phenocrysts
in Picritic Basalts from the Juan Fernandez
Islands
JAMES H. NATLAND*
ROSENSTIEL SCHOOL OF MARINE AND ATMOSPHERIC SCIENCE, UNIVERSITY OF MIAMI, MIAMI, FL 33149, USA
RECEIVED DECEMBER 10, 1998; ACCEPTED AUGUST 30, 2002
Farley et al. (1993) presented He, Sr, and Nd isotopic
data for picritic tholeiitic and alkalic basalts from the
Juan Fernandez Islands, which are located on the
Nazca plate west of Chile in the SE Pacific. Helium
isotopes were measured on gases extracted from olivine
separates crushed in vacuo, and then passed into a mass
spectrometer. This is the usual way of measuring isotopes of noble gases in subaerial ocean island basalts, as
they almost always lack the more favorable glass component characteristic of abyssal tholeiites.
There are two major islands in the Juan Fernandez
group (Fig. 1), Robinson Crusoe (also called Isla Mas a
TierraÐliterally, island nearer land) and Alexander
Selkirk (Isla Mas AfueraÐisland further away).
Because the English names, derived from Defoe's
(1719) fictional castaway and his historical inspiration
who spent some years alone in these islands, are now
official, I use them here. The islands are youthful,
with radiometric ages in the range of 40±31 Ma for
Robinson Crusoe and 13±085 Ma for Alexander
Selkirk [as summarized from various sources by Baker
et al. (1987)]. They are probably still active (Darwin,
1840). Their age progression is to the west and they are
built on Eocene crust (32±37 Ma) of the Nazca plate
(Corvalan, 1981).
The islands are remote and rarely visited by geologists. Nevertheless, the presence there of picritic basalts
uncommonly rich in olivine phenocrysts has long been
known. Bowen (1928, p. 164), after Quensel (1912),
published a photomicrograph of basalt from these
islands, unusually rich in large olivine crystals, which
he described as `having the highest proportion of normative olivine (53%) of any rock termed basalt by the
author describing it'. Johanssen (1937) gave one such
*Corresponding author: E-mail: [email protected]
Journal of Petrology 44(3) # Oxford University Press 2003; all rights
reserved.
Olivine crystals in basalt contain helium that can be extracted
for isotopic analysis. Helium-bearing olivine phenocrysts in
picritic tholeiites from the Juan Fernandez Islands, SE Pacific,
crystallized from moderately differentiated liquids. None are
xenocrysts of mantle peridotite. The helium occurs in cavities or
bubbles in inclusions best seen at high magnification in the
olivine. The olivine grew around spinel, sulfide, and bubbles that
preferentially nucleated on crystal surfaces. Inclusions within
these phases range from large cavities to tiny, faceted pits
arranged in rows. Many crystals contain curving trains of
inclusions along annealed features. The inclusions formed during
mixing between cooler differentiated magma and hotter olivinecharged magma, which accelerated vesiculation. Bubbles
nucleated on the olivine as they do on ice when stirred in
carbonated water. Mixing also induced thermal stress fracturing,
like the cracking of ice dropped into water. Cracking, irregular
extinction, and subgrain formation occurred when faceted crystals collided with each other or with conduit walls. Boundary
layer melts and bubbles were drawn quickly into the fractures.
Thus few inclusions contain equilibrium proportions of minerals
and vapor. Mantle-derived helium clearly permeated into shallow storage reservoirs, including rift zones where magmatic differentiation, mixing, and vapor exsolution were fairly extensive.
KEY WORDS:Juan
Fernandez; volatiles; inclusions; olivine; picrite
INTRODUCTION
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 3
MARCH 2003
Fig. 1. Generalized bathymetric chart of the Juan Fernandez Ridge, SE Pacific, near Chile, showing the location of principal islands and
seamounts, after Mammerickx et al. (1974), from Farley et al. (1993).
uncommonly olivine-rich basalt from the islands a
type-locality name, masafuerite. The most accessible
sources of recent geochemical data for Juan Fernandez
basalts are the studies by Gerlach et al. (1986), Baker
et al. (1987) and Farley et al. (1993). The last includes
analyses from two seamounts, one of them named
Friday after the only other principal character in
Defoe's novel.
I address the question of whether minerals containing a vapor phase with mantle-derived noble-gas isotopic characteristics actually are mantle-derived
mineral phases (xenocrysts) in their host basalts, or
whether they are phenocrysts that happened to incorporate a vapor phase carrying a mantle isotopic signature. The Juan Fernandez samples studied by Farley
et al. (1993) have 3He/4He ranging from 78 to 180 RA,
a range spanning values they attributed to mid-ocean
ridge basalt (MORB) mantle sources at the lower, and
plume-related (ocean island) sources at the higher end.
Because subaerial ocean island basalts are substantially
degassed, this question relates to the efficiency of degassing and both where in the volcanic edifice and when in
the course of magmatic differentiation degassing and
entrapment of He took place.
I have investigated several of the samples analyzed
for He, Sr, and Nd isotopes by Farley et al. (1993). I
examined olivine and associated Cr-spinel in polished
thin sections using transmitted and both regular and
interference-contrast reflected light, acquired highmagnification scanning electron micrographs of the
surfaces of magnetically separated olivine, and determined the compositions of the two mineral phases by
electron microprobe. I show that in typical Juan
Fernandez picrites of tholeiitic composition, mantle
He is contained in olivine phenocrysts, not xenocrysts,
and that most of it was evidently incorporated during
episodes of combined magma mixing and vesiculation,
the two being closely related. Late-stage post-shield
eruptives on Robinson Crusoe are basanites and these
do contain xenocrysts of four mantle mineral phases
(spinel, olivine, clinopyroxene, and orthopyroxene)
that probably contributed to their noble-gas isotopic
signature. These are not considered in this paper.
He and other noble gases are only several among
various volatile species that were present in the cavities
in the olivine, and probably not the most abundant.
Roedder (1965, 1983, 1984) pioneered the study of
fluid inclusions in olivine phenocrysts and mantle
xenoliths, and showed that CO2 is the principal volatile
constituent in them, being present both as liquid and
gas under high confining pressures. His petrographic
techniques were developed mainly to observe and conduct experiments on the fluid inclusions. The petrographic techniques used here tell nothing directly
about volatile inclusions, certainly not precisely
where they were in the minerals before the olivines
were crushed and their volatiles extracted. However,
as Roedder also observed, many inclusions may also
contain glass, minerals, and globular sulfides. These
can tell a great deal about the characteristics of
associated bubbles, even when the bubbles themselves
cannot be directly analyzed.
PETROLOGICAL BACKGROUND
Figure 2 summarizes the normative characteristics of
basalts from the Juan Fernandez Islands reported by
Baker et al. (1987) and Farley et al. (1993), and compares them with picritic Hawaiian tholeiites from
Kilauea's submarine extension, Puna Ridge (Clague
et al., 1991, 1995; Dixon et al., 1991). Tholeiitic basalts
from the two places are very similar. Investigations of
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NATLAND
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
Fig. 2. Comparison of principal normative constituents of basalts from the Juan Fernandez Islands and Kilauea±Puna Ridge. Juan
Fernandez data are from Baker et al. (1987) and Farley et al. (1993). Juan Fernandez symbols are large, as follows: Trend 1, filled gray and
black circlesÐhypersthene-normative tholeiites from Robinson Crusoe and Alexander Selkirk, with gray representing samples of this study;
Trend 2, open squaresÐnepheline-normative alkalic olivine basalts from Robinson Crusoe and Alexander Selkirk; half-filled diamonds
enclosed in fieldÐbasanites from Robinson Crusoe. Kilauea±Puna Ridge symbols (dashed trend KP) are small filled right-pointing triangles
(Clague et al., 1995). Trends 1 and 2 converge along the Ol±Di tie-line, toward the Ol apex.
volatiles in basalts from Kilauea bear directly on this
study at many points, and together with results here
point to general processes that act in the formation of
picrites.
Farley et al. (1993) noted that their Juan Fernandez
sample suite is biased toward olivine-rich samples
deliberately collected for noble-gas measurements.
Partly as a result of this sampling, two trends are
evident in Fig. 2, one extending from the olivine corner
first to more strongly hypersthene-normative or tholeiitic compositions, then ultimately to quartz tholeiite,
the other to successively more strongly nephelinenormative undersaturated alkalic compositions. The
trends diverge from a common point near the olivine
corner, among picritic basalts. There, individual samples vary only slightly in degree of silica saturation or
undersaturation. The normative distinction among the
picritic basalts may therefore be small, having significance only for more evolved rocks. Indeed, the normative distinction between the two trends stems almost
entirely from about 1±2 wt % difference in SiO2 contents, this widening with decrease in MgO contents.
Otherwise, rocks from the two trends are similarly
sodic and potassic and they have similar concentrations
and proportions of incompatible trace elements (Farley
et al., 1993). The samples studied lie along the tholeiitic
trend in Fig. 2 and encompass about two-thirds of the
range in normative hypersthene, with the exception of
one sample, a picrite that is very slightly Ne-normative.
The extended tholeiitic trend parallels that for Puna
Ridge. However, whereas the latter reaches strongly
quartz-normative compositions, the Juan Fernandez
tholeiitic trend is shifted in its entirety to the left,
lying almost exclusively in the Di±Ol±Hy ternary.
This is a consequence of somewhat higher total alkalis,
especially Na2O, and somewhat lower SiO2 and Al2O3
contents. A closer match in all these characteristics is
provided by tholeiites of the Honomanu volcanic
series of Haleakala volcano, island of Maui (see
Chen et al., 1991). The Juan Fernandez silica-undersaturated sequence is more extensively developed on
Robinson Crusoe than on Alexander Selkirk. On
Robinson Crusoe, a post-shield basanitic series followed that is even more strongly silica undersaturated (Fig. 2). The diversity of basalts sampled
on Robinson Crusoe is undoubtedly enhanced by the
greater age and more advanced stage of erosional
dissection of this volcano than of Alexander Selkirk
(Baker et al., 1987).
Projection of the tholeiitic and alkalic trends toward
the olivine apex in Fig. 2 is an indication that the
compositions of many of the basalts are olivinecontrolled. That is, the proportion of olivine phenocrysts they contain largely determines variations in
rock compositions. The olivine is accumulative, as
Bowen (1928) argued, and the bulk compositions of
the rocks therefore do not represent liquids. Indeed,
some samples reported by both Baker et al. (1987) and
Farley et al. (1993) contain 19±22% MgO, with 35±
45% of olivine in the norm.
Olivine control is especially evident on a plot of Ni vs
MgO (Fig. 3a). Ni partitions strongly into olivine, but
varies as a function of forsterite content (Hart & Davis,
1978), being lower in more iron-rich olivines. Separate
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NUMBER 3
MARCH 2003
Fig. 3. Comparison of chemical attributes of tholeiitic and alkalic olivine basalts from the Juan Fernandez Islands and Kilauea±Puna Ridge.
Symbols are as in Fig. 2. Oxides are given in weight per cent, trace elements in parts per million. (a) MgO vs Ni; (b) MgO vs total iron as FeO;
(c) MgO vs K2O; basanites are plotted on this diagram only; (d) MgO vs Zr; (e) TiO2 vs CaO. Schematic trends: 1, olivine-controlled basalts;
2, olivine±plagioclase±clinopyroxene-controlled basalts; dashed, hybrids produced by mixing. In (a), the dashed and shaded trends show
effects of control of olivine with different compositions on trends for Kilauea±Puna Ridge (KP) and Juan Fernandez (JF). In (b), lines of
constant FeOT show the effects of control of olivine having different compositions, as indicated. In (c) and (e), lines with arrows indicate
`olivine-control' trends. In (e), low-CaO Juan Fernandez tholeiites at given TiO2 and lying between Trends 1 and 2 are encircled by a
continuous line. These appear to be hybrid lavas produced by mixing between primitive, olivine-rich basalt along Trend 1, and strongly
differentiated, high-TiO2 basalt well along Trend 2.
trends for Kilauea and Juan Fernandez suggest that
the latter basalts in most cases carry more iron-rich
olivine, as does a plot of total FeO vs MgO (Fig. 3b).
Horizontal trends in Fig. 3b through the principal data
clusters for Kilauea±Puna Ridge and Juan Fernandez
indicate the nominal FeO contents of olivine of
representative compositions controlling the trends. At
Kilauea±Puna Ridge, olivine megacrysts have an
average composition of about Fo88±89; such olivine
has about 2500±2800 ppm Ni (Clague et al., 1995).
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NATLAND
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
Juan Fernandez basalts have a spectrum of higher
FeO contents, giving a range in olivine compositions
of Fo88±85; these olivines have 2000±2500 ppm Ni
(electron-microprobe measurements). The control of
olivine on Ni in rock compositions is shown in Fig. 3a.
Plots of incompatible elements vs MgO (Fig. 3c and
d) are similar, showing slightly higher concentrations
of these elements in Juan Fernandez tholeiites and
alkalic olivine basalts than in the Puna Ridge picritic
glasses. K2O contents in a few Juan Fernandez samples
are low, perhaps because of groundwater leaching.
Trends for tholeiitic glasses from Puna Ridge and for
aphyric tholeiites from Juan Fernandez have inflections departing from olivine control at about 7%
MgO, at which point at Juan Fernandez plagioclase
joins the liquidus, and clinopyroxene shortly thereafter. This is shown schematically in Fig. 3c and d,
with the effects of olivine separation on Kilauea glasses
and aphyric Juan Fernandez basalts occurring along
Trend 1, and multiphase control at lower MgO content along Trend 2. The MgO value at the inflection in
both diagrams corresponds to liquid MgL [ ˆ 100Mg/
(Mg ‡ Fe2‡)] of 60. The value of MgO (or MgL) for
this inflection among Juan Fernandez aphyric basalts
is somewhat imprecise, because fewer analyses are
available to define it than for the Puna Ridge glasses;
there is a wider range of Juan Fernandez tholeiitic and
alkalic compositions, and, as discussed next, some of
the rock compositions have been influenced by mixing.
The olivine-control trends along which each sample
lies also probably have slightly different inflection
points.
Some of the elevated concentrations of incompatible
elements in Juan Fernandez picritic tholeiites appear
to have resulted from mixing between primitive and
porphyritic olivine-rich basalts originally lying at the
high-MgO end of Trend 1 in Fig. 3c and d, and highly
differentiated basalts, lying near the end of Trend 2.
Several of the tholeiites I have studied in detail,
for example, fall along such hypothetical mixing
trends. Mixing between two such magmas would produce a hypersthene-normative hybrid, but with concentrations of alkalis and other incompatible elements
reaching those in nepheline-normative alkalic olivine
basalts. That mixing probably involved a primitive,
porphyritic end-member is indicated for some samples
especially by a plot of TiO2 vs CaO (Fig. 3e). The
effects of addition of just magnesian olivine to various
compositions along Trend 2 are indicated. However,
olivine is not a liquidus mineral along Trend 2, and
it was not simply added to those magmas, like adding salt to soup. Instead, it was added as part of a
porphyritic primitive magma slurryÐby magma
mixingÐwith the highest proportion of the slurry
being in those samples labeled `mixed', with about
2% TiO2 content and containing about 30% olivine
phenocrysts.
Wright & Fiske (1971) and Clague et al. (1995)
demonstrated that mixing between primitive and differentiated magmas was an important process in
tholeiite petrogenesis at Kilauea±Puna Ridge. However, comparison of bulk-rock and glass compositions
shows that mixing there occurred primarily between
primitive magmas and only somewhat more differentiated magmas, both lying along or very close to
the olivine-controlled portion of the differentiation
sequence. Only a few Juan Fernandez picrites
fall along Trend 1. There mixing was more often
between more strongly contrasting primitive and
differentiated magmas.
In summary, one of the most significant attributes
of the basalts from Juan Fernandez is the close
resemblance, but not identicality, to sequences of
Kilauea and Puna Ridge. Tholeiitic picrites from
both Robinson Crusoe and Alexander Selkirk have
slightly higher total alkalis, lower silica and alumina,
and somewhat greater concentrations of highly incompatible elements. The influence of olivine control, and
the point during crystallization when olivine control
ceases, are much the same. Magma mixing was also
important, but among the several Juan Fernandez
tholeiites studied, it occurred between more extreme
magma types, which enhanced alkali concentrations
among the hybrid rocks. The Juan Fernandez
volcanoes are nowhere nearly as large as any in the
Hawaiian archipelago, yet this appears to be the only
other island group built on old ocean crust in the
Pacific where low-alkali tholeiitic basalts have erupted
in any abundance. Detailed field relationships
between tholeiites and silica-undersaturated alkalic
picrites are still uncertain (Farley et al., 1993), but the
basanites, which contain upper-mantle xenocrysts,
clearly are the youngest eruptives on Robinson Crusoe
(Baker et al., 1987). In this respect, they are comparable to the xenolith-bearing undersaturated lavas of the
post-shield rejuvenated, once termed post-erosional,
stage of Hawaiian volcanism.
The mixing pattern of Juan Fernandez tholeiites is
apparent in the compositions of olivine and spinel,
as discussed below, and is important to the
mechanism of incorporation of volatiles into olivine
that I propose.
ANALYTICAL PROCEDURES
Polished thin sections were examined using both transmitted and reflected light at magnifications from 63
to 1600. Best results in reflected light were obtained
using differential interference contrast, which highlighted residual relief from polishing around both
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JOURNAL OF PETROLOGY
VOLUME 44
mineral inclusions and bubbles. Oil immersion at high
magnification heightened reflectivity contrasts between
iron±titanium oxide minerals and spinel, and reveals
myriad minute sulfides. The study is concentrated on
the phenocrysts themselves, the shapes of their inclusions, and the other phases besides volatile constituents
those inclusions contain (minerals and/or glass). This is
dictated by petrographic technique. Solid phases are
the only ones that can be seen on polished or broken
surfaces. Accordingly, I identify varieties of cavities,
assuming that these were the vessels that once contained volatiles, whether fluid or vapor, assuming
that rupture of these cavities released all original volatiles, including the noble gases that were measured for
isotopes. In discussion below, cavities that lack mineral
phases and either glass or devitrified glass are termed
`empty', although they originally contained volatiles.
Individual olivine grains originally concentrated by
magnetic separator as splits for the isotopic analyses
were mounted on stubs, coated with Au±Pd, and
studied by scanning electron microscopy (SEM) at
various magnifications at the Rosenstiel School,
University of Miami.
Photomicrographs were scanned at 300 dpi into a
computer and then scaled and adjusted for contrast
using Adobe Photoshop. In some cases, enhancements
were applied to selected portions of photomicrographs,
or to individual objects on them.
Mineral compositions were determined on six tholeiites, two of them from Robinson Crusoe and four
from Alexander Selkirk, selected from the samples analyzed for major oxides by Farley et al. (1993). One
additional sample from Alexander Selkirk was studied
by electron microprobe, but it has no rock analysis.
On the basis of mineralogy and petrography, it is a
tholeiite. Three additional samples, one of them
analyzed, from Alexander Selkirk were examined in
thin section only. Three basanites were also studied
by electron microprobe, but the data are not reported
here. Mineral analyses were normalized to the same
mineral standards in all cases. The standard for olivine
was US National Museum (USNM) San Carlos olivine. For spinel, USNM New Caledonia chromite was
used except for TiO2, total iron as FeO, and MnO,
which were normalized to the standard values for
USNM Tiebaghi ilmenite.
PETROGRAPHIC SUMMARY
In addition to containing olivine and rare skeletal
plagioclase phenocrysts, tholeiitic picrites from the
Juan Fernandez Islands have fine- to medium-grained
groundmasses consisting at least of plagioclase, clinopyroxene, and either titanomagnetite or ilmenite
NUMBER 3
MARCH 2003
and titanomagnetite. The phenocryst±groundmass
sequence gives the order of crystallization in all samples
regardless of grain size, based on idiomorphic relationships. This differs slightly but systematically from
Kilauea tholeiites in that plagioclase in all cases precedes clinopyroxene instead of joining the liquidus at
the same time or shortly after (see Wright & Fiske,
1971; Helz & Thornber, 1987; Clague et al., 1995).
The distinction does not hold for melt inclusions in
olivine phenocrysts, as described below.
The lavas are variably vesicular. The more crystalline ones, which presumably came from flow interiors,
have scattered, round, pinhole-sized vesicles. The more
obviously quenched ones from flow tops have a greater
percentage of vesicles that are irregular in shape and
up to 3 mm in longest dimension. Flow-top plagioclases
are small, acicular, and separated by a dark cryptocrystalline mesostasis; flow-interior plagioclases are
larger, usually tabular, and are generally separated
by more coarsely crystallized clinopyroxene and oxide
minerals. Oxide minerals vary greatly in morphology,
tending to be extremely elongate with elaborate cellular skeletal and dendritic morphologies in quickly crystallized flow tops, and more nearly equant but far less
cellular and skeletal in flow interiors. In flow interiors
ilmenite tends to elongate rhombic cross-sections and
titanomagnetite to equant morphologies bounded by
silicate crystal margins. Neither orthopyroxene nor
primary sulfide occurs in the groundmass of any
sample.
The rock groundmasses indicate stages of differentiation that are nearly independent of bulk MgO contents; the latter depend more on the circumstantial
extent of olivine accumulation. In less differentiated
groundmasses, skeletal olivine is intergrown with euhedral plagioclase and encloses minute Cr-spinel. More
differentiated groundmasses have no skeletal olivine.
Tabular or microlitic plagioclase predominates, and
it is partly encased by clinopyroxene. Two such samples also have fairly large anhedral clots of intergrown
ilmenite and titanomagnetite, in addition to much
smaller crystals of both minerals scattered throughout
the groundmass. Some of the clots are attached to
rims of olivine phenocrysts. Another sample has
a small, felsic, and probably andesitic clot lacking
clinopyroxene, and having only sparse and very tiny
oxide minerals. Several of the rocks have swirls of
different color, crystallinity, and vesicularity. Oxide
minerals have different morphologies, being skeletal
in the coarser-grained portions of the rocks, and both
smaller and more equant in the finer-grained portions.
The darker colors and contrasting crystal morphologies are probably an indication of a higher proportion
of oxide minerals and clinopyroxene, and also different
crystallinity, crystal morphology, and grain size. The
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NATLAND
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
swirls therefore probably result from compositional
heterogeneity and are a consequence of incomplete
magma mixing.
Among these lavas, the contrast between those with
olivine crystallizing in the groundmass and those lacking olivine in the groundmass probably reflects a contrast in the proportions and compositions of liquids
that originally crystallized and accumulated olivine,
and those that now host it, an indication of mixing
between primitive and differentiated magmas. The
extreme contrast is the rock in which olivine phenocrysts enclosing Cr-spinel are intergrown with ilmenite
and magnetite at their rims. A secondary contrast has
to do with the proportion of ilmenite to titanomagnetite, and even the absence of ilmenite in some samples.
This may be related to different Fe/Ti ratios in the
crystallizing liquids (Wright & Fiske, 1971).
Olivine phenocrysts are present in various proportions (up to 28%) and sizes; many are rather large
(09 cm). In general, these conform to the five classes
identified in tholeiites of the 1959 eruption of Kilauea
volcano by Helz (1987), namely: (1) large, irregular
blocky grains, often with planar extinction discontinuities, or kink bands; (2) euhedral and skeletal grains;
(3) resorbed grains with original morphologies not
evident; (4) angular to conchoidal fragments; (5)
grains with swarms of inclusions of opaque minerals
and pale glass. However, I do not attach the same
genetic significance to the classification as did Helz
(1987), nor are the distinctions between the classes
necessarily sharp.
For example, we may compare two examples of
blocky, irregular crystals (Fig. 4a and b) with two
euhedral crystals (Fig. 4c and d). The olivine in
Fig. 4a is faceted on every side, even though it is
irregular in shape. The olivine in Fig. 4b is not faceted
only because of small chips and indentations in its
outline. Otherwise, it would be almost identical to the
olivine of Fig. 4d. Almost all of the `blocky and irregular' olivine in our samples is like this. The original
outlines have simply been chipped away by some
mechanical process, but in many cases are still partly
present. Thus, the faceted, or nearly faceted, aspect of
both classes is their most important feature, with the
precise shape depending mainly on the cross-section of
the mineral, or of intergrown grains, in the slide, less
whatever has been nibbled from the edges. Therefore
no fundamental distinction exists between the two
classes, the irregular grains are not xenocrysts, nor
did they necessarily crystallize in some other magma
than euhedra of the same general size in the same
sample. Both types clearly crystallized while being surrounded by melt. As to broken grains (Class 4), these
typically conform in shape to the fracture-bounded
internal segments of any of the olivine crystals shown
in Fig. 4a±e, and thus more than likely originally were
parts of phenocrysts as well. There are very few strikingly skeletal grains (Class 2) or those with amoeboidal
outlines (Class 3). The latter are simply unusual sections through strongly skeletal crystals rather than the
result of resorption.
A few grains are combined in two- to three-crystal
aggregates. Some grains are rounded, or have rounded
interiors surrounded by a normally zoned rim with a
faceted margin. In the most porphyritic samples, about
10% of the Class 1 and Class 2 olivine crystals exhibit
coarse sub-grain development, irregular extinction, or
kink banding (Fig. 4e and f). Such grains, and rounded
grains, may be common in one sample but absent in
another. One sample contains a dunitic xenolith about
1 cm in diameter.
Helz (1987) suggested that isolated Class 1 crystals in
basalts from Kilauea `were produced by disaggregation of coarse, mildly deformed dunite' (p. 712).
Among these grains are those with planar deformation
surfaces such as kink bands and rectangular subgrains.
Helz (1987) cited experimental data (Raleigh, 1968;
Carter & Ave Lallement, 1970; Raleigh & Kirby,
1970) to the effect that such deformed olivine forms in
a crystalline matrix subject to non-hydrostatic stress.
She concluded, `kink-banded crystals surrounded by
melt must have been deformed in a very different
environment' (p. 712). Partly from this suggestion,
Clague & Denlinger (1994) argued that Class 1 olivine
in Kilauea tholeiites, particularly that with kink banding, was probably scavenged from dunitic cumulates
residing deep within the volcano.
The faceted, euhedral but deformed crystals such as
those in Fig. 4e and f, were not scavenged from dunite.
Dunite is a cumulate, indeed an adcumulate, in which
an original loose packing of minerals has been transformed by a process of grain boundary dissolution,
reprecipitation, and textural equilibration (Hunter,
1987, 1998) to a collection of intergrown grains almost
none of which retain their original euhedral morphology (Fig. 4g). The presence of subgrains in faceted
olivine means instead that euhedral phenocrysts can
indeed experience non-hydrostatic forces within a
magma; that their edges may become chipped or even
rounded by mechanical abrasion; and that some of the
grains can break into fracture-bounded segments
during magma transport and flow. Later I suggest a
mechanism for this that also has bearing on the origin
of certain types of melt and fluid inclusions.
In all samples, olivine typically hosts one to several
much smaller (503 mm 02 mm, and typically
01 mm 007 mm), euhedral, black crystals of chromian spinel. Rarely, some dozens of spinel grains and
other inclusions occur in a single grain. These correspond to Class 5 olivines of Helz (1987), but rather
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Fig. 4. Features of olivine phenocrysts. Olivine compositions (Fo) are indicated. (a) Blocky, euhedral olivine with almost no inclusions,
sample PIN-8, Robinson Crusoe; transmitted light. (b) Block, irregular olivine enclosing some spinel. Sample MF-C4, Alexander Selkirk.
Lines are drawn to indicate the original euhedral shape of the crystal. Transmitted light. (c) Euhedral olivine phenocryst enclosing a few
spinels, and containing numerous ellipsoidal cavities parallel to crystal faces at upper right and lower left. The lower left face is partly broken.
Sample MF-C4, Alexander Selkirk. Transmitted light. (d) Euhedral olivine similar in shape to that in (b), but with a partly broken edge at
lower right. Sample PIN-12, Robinson Crusoe. Crossed nicols. (e) Euhedral olivine with subgrains. Sample PIN-12, Robinson Crusoe.
Crossed nicols. (f) Two-crystal aggregate, having partly faceted and partly broken edges, and containing parallel deformation lamellae with
irregular extinction. Sample PIN-12, Robinson Crusoe. Crossed nicols. (g) Dunite cumulateÐa xenolith from Ta'u Island, American Samoa.
Absence of euhedral outlines, and the olivine with faint subgrains at lower right, should be noted. Many crystals have 120 triple junctions,
indicating textural equilibrium in a monomineralic cumulate (Hunter, 1996). Sample 82MT-X2. Crossed nicols.
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HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
than considering such grains as a distinctive category,
they are only one extreme in a population of olivine
grains having from `none' to `many' such inclusions. A
few spinel grains occur at olivine rims where they may
be skeletal. In samples having the less differentiated
groundmass, spinel also occurs as euhedral or skeletal
isolated phenocrysts, generally smaller than spinel
within olivine. Some rim and groundmass spinel has
more reflective titanomagnetite margins; in the basalt
with most differentiated groundmass, the rim spinellid
mineral is actually chromian titanomagnetite intergrown with ilmenite. In the two samples with the
most olivine phenocrysts, a few olivine grains enclose
spherical sulfides either as isolated or multiphase inclusions. The larger sulfides comprise both pentlandite
[(Fe, Ni)9S8] and pyrrhotite. One other more differentiated sample has a trace of sulfide in the inclusions in
one part of one olivine grain. Absence of sulfides in
phenocrysts of most samples suggests that their S was
degassed before the olivine crystallized. In all samples,
degassing reduced S concentrations to well below sulfide saturation before crystallization of the groundmass.
PETROGRAPHY OF INCLUSIONS
Photomicrographs and scanning electron micrographs
are presented generally in the order of increasing
degree of magnification, but are to some degree interspersed to illustrate aspects of the same phenomena at
different magnifications.
General characteristics of inclusions in
olivine phenocrysts
As mentioned above, olivine phenocrysts have varying
percentages of mineral±melt±vapor inclusions, ranging within the same sample from none to many
(Fig. 4a; Fig. 5a±c). The olivine ranges from blocky
and euhedral (Fig. 4a, c, and d) to fairly rounded
(Fig. 5a±c). In transmitted light, both melt inclusions
and spinel appear as black inclusions in photomicrographs. Many grains have central cores that
are fairly densely charged with inclusions, and local
concentrations of spinel-bearing inclusions elsewhere.
The grains in Fig. 5a±c illustrate a common modality
of inclusion arrangement, that of linear or curving
arrays. These are the features that Roedder (1965)
originally described as secondary inclusions, and that
he interpreted as annealed fractures along which inclusion material was brought into the crystal. In some
cases these are parallel. In Fig. 5a, three linear arrays
(near arrows) converge on a point outside the boundary of the present olivine. In all the olivine, such arrays
do not correspond to the current generation of
fractures crossing the grains. The olivine of Fig. 5c
combines a dispersed arrangement of inclusions with
curvilinear arrays.
The olivine of Fig. 5d is distinctive within the sample
in having an extraordinarily dense concentration of
small inclusions and some large ones carrying spinel.
Some rows of spinel parallel the upper and lower crystal boundaries, and there is one large and rounded melt
inclusion now crystallized to spherulitic basalt at center left. This grain has a younger generation of fractures slightly oblique to the linear arrays of spinel and
the upper crystal faces. The inclusion-bearing region is
bounded by a fairly wide rim, free of inclusions, but not
of uniform width. At high magnification in reflected
light (Fig. 5e) the dense array of small inclusions is seen
to consist of elliptical to rectilinear cavities in the olivine; these are variously filled with devitrified melt
(dark gray), spinel (light gray) and sulfide (bright).
Dark gray shadowed pits in some of these are bubbles.
The larger of these are oriented parallel to the linear
arrays of large spinels in Fig. 5d.
Although the orientation of these inclusions exhibits
a strong crystallographic control, there is no indication
that they were incorporated into the olivine along
fractures. This particular olivine corresponds to perhaps the extreme variant of class 5 (grains that contain
a very high density of tiny inclusions) in the classification of Helz (1987), even to the extent of containing
sulfides. She noted the rarity of such grains in the 1959
Kilauea eruptives and in samples from the Kilauea
Iki lava lake. Schwindinger & Anderson (1989) and
Clague et al. (1995) illustrated similar olivine densely
charged with inclusions from Hawaii.
Farley & Craig (1994) found highly variable yields of
Ar in several splits of olivine grains from the same
sample (PIN 12; Robinson Crusoe). This apparently
reflects the existence of variable inclusion densities in
different olivine crystals seen during petrographic
examination.
Inclusion arrays and indications of
deformation
Figure 6a±c shows several olivine grains containing
both inclusions and subgrains or planar extinction
discontinuities bounding kink bands. Usually there
is no relationship between curvilinear arrays of inclusions and subgrain or kink boundaries (Fig. 6a, near
arrows, and Fig. 6c). In Fig. 6c, the rounded and
deformed interior of the olivine is carapaced with an
undeformed outer zone. Three trains of inclusions converge toward the upper right portion of the crystal,
reaching the outer zoned margin. The mineral was
first deformed, then incorporated into magma, then
rounded, and then the curvilinear arrays of inclusions
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Fig. 5. Features of inclusions in olivine phenocrysts photographed in transmitted light. Olivine compositions (Fo) are indicated. (a)±(c) show
olivine phenocrysts in sample MF-C2, Alexander Selkirk. (a) Olivine phenocryst with several linear inclusion arrays trending toward the
upper right. (b) Olivine phenocryst with nearly orthogonal inclusion arrays. (c) Rounded olivine phenocryst with many inclusion arrays.
(d) and (e) show sample MF-S1, Alexander Selkirk. (d) An olivine phenocryst with a core densely charged with inclusions, and a mantle with
no inclusions. (e) Detail of a portion of the same olivine phenocryst, photographed using differential interference-contrast reflected light. In the
inclusions, which are recessed by polishing, dark gray is glass, light gray is spinel, and bright is sulfide.
formed. Figure 6b provides an example in which several spinel-bearing inclusions lie along a planar deformation boundary (near line with arrowheads). In this
case, the deformation occurred after incorporation
of the inclusions, and the location of the planar
deformation boundary may have been guided by internal stresses in the olivine associated with the inclusions.
Incorporation of inclusions along curvilinear arrays
proceeded even into the eruptive history of some of the
rocks. Some olivine is decorated with triangular dendrites in optical continuity with the mineral interiors
(Fig. 6d). Growth of such dendrites rather than
continued uniform growth of the exterior of the crystals
indicates subjection to suddenly higher cooling rates,
or undercooling (Donaldson, 1976). In these examples,
formation of dendrites was the last growth the
olivine experienced before fairly rapid cooling of the
groundmass following eruption. Scanning electron
micrographs show that the dendrites have the typical
faceted pyramidal terminations of single olivine crystals (Fig. 6e and f ), and that they budded preferentially on the curving edges of flat, stepped surfaces that
are bounded by fractures orthogonal to the surface of
the olivines (Fig. 6e). In detail the orthogonal fracture
surfaces form sets of fractures with small offsets, but
curving in total 90 . The stepped morphology of the
crystal edge is seen in thin sections to be particularly
common on crystal faces paralleling the c-axis in rocks
with a fine-grained groundmass, in which it is evidently
a consequence of undercooling. The same crystals
usually have skeletal projections at terminations on
101 or 021 crystal faces. The growing olivine was first
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HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
Fig. 6. Deformed olivine and olivine with epitaxial dendritic overgrowths. Olivine compositions (Fo) are indicated. (a) Olivine phenocryst
with subgrains. Crossed nicols. Sample MF-S1, Alexander Selkirk. (b) Kink-banded olivine in the same sample. Crossed nicols. (c) Olivine
with irregular extinction. Same olivine as in Fig. 4c. Crossed nicols. Sample MF-C2, Alexander Selkirk. (d) Olivine phenocryst with a dendritic
rim and an inclusion train crossing one dendritic projection. Transmitted light. Sample MF-C2, Alexander Selkirk. (e) Scanning electron
micrograph of dendritic projections growing on a stepped olivine crystal surface. Sample MF-3, Alexander Selkirk. (f) Scanning electron
micrograph showing detail of olivine elsewhere on the same olivine.
subjected to stresses that allowed the orthogonal fractures to form, defining the small stepped offsets on
the mineral surface. At still higher undercooling, the
dendrite buds then nucleated preferentially at the
locations of strongest curvature, perhaps because this
is where surface tension of the growing mineral surface
was most disrupted.
The olivine with a dendritic surface in Fig. 6d is
complexly fractured, and also has a number of linear
or curvilinear trains of inclusions, some of them fairly
obviously arranged along fractures, some of them not.
One of these arrays, however, is along a fracture (near
line with arrowheads) cutting through the pyramidal
cross-section of one of the dendrite studs on the mineral
boundary. This is shown enlarged in the inset. The
inclusion array thus formed after the dendrite began
growing, that is, when cooling rates abruptly increased
sometime just before eruption, or perhaps even while
the lava was flowing. In one sample, an inclusion array
formed late in the groundmass, within a clinopyroxene
that subophitically encloses tabular plagioclase and
that is adjacent to an olivine phenocryst.
Varieties of inclusions
To this point I have mainly described curvilinear
arrays of inclusions that apparently define fracture
surfaces, or annealed fracture surfaces, in the olivine.
Several other varieties occur, including cavernous
inclusions, mineral±melt±bubble inclusions, open cavities around spinel, and faceted sub-microscopic pits,
the last of which can be seen only at very high
magnification.
Cavernous and ellipsoidal inclusions
A tholeiite (sample MF-C4), contains numerous
olivines with large cavernous and ellipsoidal cavities.
These are especially evident in reflected light (Fig. 7)
and occur in both rounded (Fig. 7) and euhedral
faceted grains (Fig. 4b). Some of the cavities have
angular surfaces, but most are ellipsoidal, with their
long axes parallel to the long dimension of the mineral,
which is the crystallographic c-axis (Fig. 8a). Some
of the cavities clearly acted as stress guides for later
contraction fractures, one set of which in Fig. 7 is
prominently nearly concentric to the outer surface of
the mineral. Two spinel inclusions in the olivine also
contain such cavities. One linear array of microinclusions diagonally cuts the lower third of the
mineral (double-headed arrow).
Scanning electron micrographs reveal the tendency
of the larger, angular cavities in the olivine from this
sample to have faceted morphologies (Fig. 8b±d).
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Fig. 7. Composite photomicrograph in reflected light of an olivine phenocryst with large, ellipsoidal, cavities (black), and which also encloses
several Cr-spinels (white). Light gray spots are cavities below the polished surface of the grain revealed by internal reflections. The linear
inclusion train to the right of the arrowed line should be noted. Sample MF-C4, Alexander Selkirk.
Some smaller cavities contain euhedral crystals of
plagioclase and pyroxene (Fig. 8c and d). Another
sample, MF-C3, has olivine with faceted inclusions
partly filled with devitrified melt (Fig. 8e). A spinel
crystal once occupied the sharply curving central cavity in this inclusion. Its space is bounded at the top by a
bubble. Other inclusions in olivine from this sample,
although exquisitely faceted, are empty (Fig. 8f ).
Ellipsoidal inclusions in sample PIN-12 are teardrop
shaped (Fig. 8g). Orientation along a common fracture
is suggested. The inclusion evidently was once larger,
being surrounded now by a zone of secondary olivine
growth, as identified by dispersed X-ray emission
analysis.
Mineral±melt±bubble inclusions
The most easily seen variety of inclusion in Juan
Fernandez olivine is the large, usually isolated cavity
containing devitrified melt, crystals of silicate or spinel,
and one or more bubbles. In two samples, PIN-12 and
MF-S1, these inclusions also contain sulfide. Such
multiphase inclusions range in size down to the densely
swarmed dust-sized inclusions in the olivine of
Fig. 5d. Reflected light reveals dendritic herringbone
clinopyroxene (Fig. 9a and b), sulfide and spinel. In
Fig. 9a, the open cavity or bubble is in the spinel; in
Fig. 9c, the bubble, with a sulfide at its rim, is in the
devitrified melt (dark gray). In Fig. 9a and c, the
mineral that resembles a stack of blocks is dendritic
clinopyroxene. The dendritic morphology of the pyroxene is developed spectacularly in Fig. 9b, where its
appearance is heightened in relief by alteration of the
surrounding glass to soft clay minerals.
Sulfides show up prominently in multiphase inclusions of sample PIN-12 (Fig. 9a). Some cluster on the
surface of the large spinel at the top. Oil immersion
allows optical resolution of sulfides both on the
polished surface of the section and, by nearly matching
the index of basaltic glass, to some extent within it, by
internal reflection off surfaces of sulfide globules. These
grow fainter with depth over the distance of a few
microns. The effect is better illustrated by a direct
comparison of the same inclusion, provided in Fig. 9d
and e. Both photomicrographs are in reflected light,
but the larger of the two, using oil immersion, clearly
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NATLAND
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
Fig. 8. Features of large cavities in olivines. (a) Close-up of some cavities from the olivine of Fig. 6, reflected light. (b) Scanning electron
micrograph of a cavernous faceted cavity in another olivine of sample MF-C4. (c) Small faceted cavity with intergrown crystals of pyroxene
and plagioclase. Sample MF-C4, Alexander Selkirk. (d) Faceted cavity in an olivine phenocryst containing glass, the rounded mold of a
Cr-spinel, and a bubble at top right. Sample MF-3. (e) Small rounded cavity with a large tabular crystal of plagioclase and cellular
clinopyroxene tending to branch. Sample MF-C4. (f) Regularly faceted sawtooth cavity. Sample MF-3. (g) Ellipsoidal cavity along a
fracture. Sample PIN-12, Robinson Crusoe. The smooth material lining the cavity is olivine, as determined by peak heights using energydispersive X-ray spectrometry.
highlights the tiny sulfide droplets that are dispersed
in the glass, and diminishes the reflectivity contrast
between silicate minerals and glass. The `scorpion's
hook' here is clinopyroxene. There is a small bubble
in this inclusion, at the upper left in Fig. 9d.
Most multiphase inclusions contain glass. In
some cases they have a great deal of glass, even
when host olivine is in fairly coarse-grained rocks.
Clinopyroxene in the inclusions is spherulitic to
cellular±dendritic: morphologies typical of this mineral
in rapidly cooled basalts. With glass present, the
morphologies resemble those commonly seen within
1±3 cm of the glass rims of submarine pillow basalts
(see Bryan, 1972; Natland, 1979). They suggest undercoolings of tens of degrees to perhaps more than 100
(Kirkpatrick, 1979). The other consequence of the
presence of glass is that the bubbles in them are actually trapped vesicles that were quenched in place, just
as they are in pillow basalts. They are not bubbles that
nucleated and grew in situ, as a consequence either of
pressure reduction during eruption or of crystal growth
within the inclusion during slower cooling of the host
groundmass. Then there would have been no glass.
These inclusions demonstrate that olivine occluded
spinel, and that inclusions also often contain partly
devitrified glass and a vapor phase plus sulfide in a
minority of samples. Considered individually, many
inclusions contain disproportionately high percentages of spinel or spinel and sulfide. As first noted by
Roedder (1965), olivine phenocrysts are clearly the
preferential surfaces of nucleation at least for spinels
and bubbles, often together. The tendency was
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Fig. 9. Photomicrographs of multiphase inclusions. (a) Inclusion containing dark gray glass, light gray dendritic clinopyroxene, a large white
spinel, and dozens of tiny, bright sulfides. Oil-immersion reflected light. Sample MF-S1. (b) Spectacular pyroxene dendrites, with relief
enhanced by alteration of surrounding glass, in an inclusion. Sample MF-C4. (c) Several small multiphase inclusions, two with dendritic
clinopyroxene and spinel, the larger with a bubble, in normal vertically illuminated reflected light. Sample PIN-8. (d) and (e) are two views of
the same multiphase inclusion, with an unusual hook-like faceted and partly cellular aggregate of clinopyroxene, two marginal bubbles, and
numerous sulfides. Sample MF-SI. Many more sulfides show up under oil-immersion reflected light (e) than under standard vertical
illumination (c). (f ) Skeletal spinel at the margin of an olivine phenocryst. Sample MF-C4. Standard vertical illumination.
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HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
undiminished by high cooling rates experienced near
the point of eruption, when dendritic outer zones of
olivines grew; those zones still incorporated skeletal
spinel (Fig. 9f ).
The disproportionate share of the volume of inclusions belonging to spinel and sulfides is best seen in
smaller inclusions, whether they formed along curvilinear arrays or in dense swarms like those of Fig. 5e.
Figure 10a shows a pitted spinel occupying almost all
the space of one inclusion, but there is a small amount
of devitrified melt, and a bubble, at the upper left.
Figure 10b±d shows spherical sulfide inclusions, some
with a meniscus separating the sulfide, which was an
immiscible liquid phase, from a vapor phase. The olivine grew completely around the globular sulfides and
their adjacent volatile fluids. One inclusion in Fig. 10c
is almost completely occupied by two phases, sulfide
and spinel, which divide it about equally. The inclusion to the upper left of this one is about two-thirds
filled with devitrified glass, but about one-quarter of
that is occupied by a sulfide. A bubble comprises about
one-third of the area of the inclusion. The menisci in
adjacent inclusions in Fig. 9b have different orientations (top, bottom, side) suggesting either that the
olivine was still moving and that the bubbles were
present when sulfide crystallization (at 700 C), took
place, or that the effects of surface tension and viscosity
within the inclusions prevented bubbles adjacent to
menisci from responding to gravity.
In the extreme, small inclusions can be occupied
exclusively by spinel and a bubble cavity, or sulfide
separated from a bubble cavity by a meniscus. Portions
of curvilinear arrays sometimes have only spinel
(Fig. 9e and f ); others contain only sulfide and bubbles
(Fig. 9d). Some linear arrays with spinel are crossed by
other arrays of tinier inclusions containing sulfide but
no spinel. In most cases the odd somewhat larger inclusion in the array will contain some devitrified melt with
very fine-grained dendritic pyroxene. In these rocks
there evidently were effects of physical fractionation,
first of molten silicate away from refractory spinel, then
of molten sulfide and volatile fluids away from silicate
melt, as injected material squirted its way along fractures propagating through the olivine. In the olivine of
Fig. 5d, which swarms with small inclusions, the distribution of spinel is not uniform; they are especially
concentrated just adjacent to the outer, inclusion-free
zoned rim (Fig. 10g).
Cavities adjacent to spinel
Although empty spaces can be seen around many
included spinels in reflected light, these features are
best revealed on the broken surfaces of the olivines by
SEM. Perfectly euhedral small spinel (10 mm) usually
has small, curving cavities next to one or more crystal
faces (Fig. 11a±c). The cavities exist even where the
spinel protrudes into larger rounded cavities (Fig. 11b).
Adjacent olivine also usually has stellate arrays of
short fractures, each tending to originate from a spinel
crystal interface (Fig. 11a±c). In cross-section, these
typically form flat, curving trajectories around the spinel (Fig. 11d). A view of one half of a cavity, with the
spinel removed (Fig. 11e), shows that the spinel was
attached to one flat surface, and was contained in an
otherwise spherical space now criss-crossed by fractures.
Larger spinel is usually more thoroughly embedded in
host olivines, but even these crystals are partly surrounded by empty space (Fig. 11f ). Sense of scale is
important. Under a microscope, the small spinel crystals of Fig. 11a±e would be barely visible in the individual olivine grains shown in Fig. 5. The empty spaces
next to them would not be visible at all.
Faceted, sub-microscopic pits
A considerable surprise during examinination by
SEM was the discovery of myriads of extremely small,
faceted pits on the surfaces of some olivine. In Fig. 12a,
the tiny cavities bear some relationship to larger,
cavernous cavities, as they decorate one face of a crystal
termination, perhaps 021 or 010, whereas the several
largest cavities are on the opposing and otherwise
undecorated crystal face, and some intermediate-sized
cavities are on the crystal edge where the two surfaces
meet. At the lower left, the field of small cavities
abruptly terminates at a curving string of intermediatesized cavities just above another intersection of planar
crystal surfaces. The curving string of cavities may be a
zonation boundary.
The small cavities in this cross-section are of the
order of 1 mm wide or less, and 1±7 mm long, having
cross-sections of individual or connected parallelograms with faces oriented in the direction of opposing
planar surfaces in Fig. 12a. In Fig. 12b and c, they
are closely and regularly spaced. They would be
impossible to see in transmitted light. Only the largest
might be visible in reflected light, but their shapes
would not be evident.
Similar cavities viewed in a different orientation
are shown in Fig. 12d±f. Here, the alignment and
regular orientation of the inclusions again are
evident, although a false curvature is imparted by
the conchoidally fractured surface of the olivine
being examined. The field shown in Fig. 12b is a
continuation downward on the same curving surface
as in Fig. 12e. The cavities are more coalesced and
widely spaced. Figure 12f is a larger magnification of
a portion of the field in Fig. 12d. The faceted aspect
of the cavities is spectacularly expressed (compare with
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Fig. 10. Spinel and sulfides in multiphase inclusions. (a) A rounded Cr-spinel, itself containing small inclusions recessed by polishing, occupies
most of a larger inclusion, other phases of which, including a bubble, can be seen at the upper left. Shadowing of the small inclusions in the
spinel, using differential interference contrast reflected light, reveals that they also contain bubbles. (b) Sulfides in small round inclusions along
a curvilinear array. Several of the inclusions also have menisci in various orientations. Sample PIN-12. (c) Three small inclusions along a
linear array in an olivine, each with different proportions of melt plus dendritic pyroxene, bubbles, spinel (gray) and sulfide (bright).
Differential interference-contrast reflected light. Sample MF-C4. (d) High-magnification detail of sulfide inclusions, in oil-immersion reflected
light, showing the presence of a meniscus and two sulfide phases, pentlandite (bright) and pyrrhotite (light gray) separated by an irregular
boundary. Sample PIN-12. (e) A curving array of inclusions within an olivine. Sample MF-C2, Alexander Selkirk. The inclusions are filled
mainly with spinel. The train of inclusions near the arrowed line is blown up in (f), in reflected light. The cavities near some of the spinel,
and two small isolated cavities at upper right, should be noted. (g) A swirl of tiny spinel grains near the edge of the inclusion-rich olivine of
Fig. 4f±h. Sample MF-S1, Alexander Selkirk, in differential interference-contrast reflected light.
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HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
Fig. 11. Scanning electron micrographs of spinel in olivine from sample PIN-12. (a)±(c) show relationships between small spinel euhedra,
adjacent cavities, and fractures. The growth pattern of olivine next to the rounded cavity surface in (a) suggests inward growth or reaction of
the olivine toward the spinel. (d) Face-on view of a fracture around a spinel grain. (e) A cavity like those of (a)±(d) from which the spinel has
been removed. The spinel adhered to the olivine on the flat surface at lower left. The rest of the cavity surrounding the cavity, however, was
nearly spherical. It is crossed by a number of fractures that formed by contraction of the olivine around the spinel. (f) Edge of a large spinel with
irregular space between it and surrounding olivine. The smoothly crystalline olivine next to the spinel may have resulted from reaction or
partial re-equilibration of the spinel with melt in the inclusion, and the host olivine.
Fig. 8b, e and f ). The coalesced cavities are the
inverse shape of so-called hopper crystals (Donaldson,
1976), with widened central portions and spiky projections from cavity terminations along crystal planes.
To the extent that these examples reveal, the faceted
pits are empty; they contain no spinel, no sulfide, and
no trapped glass. The olivine of Fig. 5d and e, with its
swarms of tiny multiphase inclusions, is as close to an
exception as may exist, but many of those inclusions
are larger, and they are not faceted. Whether the tiny
faceted cavities are common or fairly rare is difficult to
say without a more systematic SEM survey. They may
be confined to or at least more prevalent on crystal
terminations than in the interiors of individual grains.
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Fig. 12. Scanning electron micrographs of tiny faceted cavities in olivines, sample KF-19. (a) Tiny pits along one crystal surface, concentrated
both along one crystal edge and against an apparent external zone of the crystal devoid of cavities at lower left. Several much larger faceted
cavities are on the opposing crystal face. (b) Detail of small cavities in (a), showing their regular spacing and the tendency of faceted surfaces to
be oblique to the overall pattern of spacing. (c) Another view of closely spaced and regularly spaced tiny cavities showing a similar obliquity of
faceted surfaces to the alignment of the cavities. (d)±(g) show a series of micrographs of a single array of tiny, faceted cavities on a broken and
curving olivine surface. (d) shows the overall spacing and arrangement of cavities. (e) shows that the pattern of spacing and coalescence of
cavities changes on the mineral surface below the area of (d). (f) A close-up view of a portion of (d). In this view, the faceted cavities are more
symmetrically disposed along the general orientation of the inclusions. The contrast with (b) and (c) may be a consequence of a less oblique
angle of the mineral surface to the cavities.
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The illustrations here suggest a total porosity of about
5±10%. If submicroscopic pits are common, they
may represent a significant proportion of the volume
of volatiles trapped by olivines. In most samples, however, they are probably subordinate in volume to
cavernous cavities or those adjacent to spinel.
Tiny cavities termed micropores have been observed
in scanning electron micrographs of alkali feldspars
in granitic rocks (Walker et al., 1995). These are
produced during the subsolidus transition between
orthoclase and microcline, apparently by dissolution
into late magmatic and hydrothermal fluids. The
majority of them appear to form at temperatures
5400 C as a consequence of penetrative alteration of
their plutonic hosts. Although there is usually a clear
crystallographic control on their orientation, they are
not arranged in rows and are rarely faceted. Usually,
they are at least partially rounded and occur as spaces
between subgrains in microperthites. The only parallel
between them and the tiny, faceted cavities in
Juan Fernandez olivine is that both formed in the
presence of fluids. The latter, however, resulted from
crystal±fluid interaction at the magmatic stage; there is
no indication that the olivine surrounding the faceted
cavities is altered at all.
ORIGIN OF INCLUSIONS DURING
CRYSTAL GROWTH
A somewhat surprising impression is that olivine in
these basalts is full of holes. The occasional low total
for an olivine electron-microprobe analysis may be
quite valid. Variations in Cr2O3 contents beyond
detection limits may have nothing to do with the olivine crystal structure; instead, the microprobe beam
probably encountered some very tiny spinel. Concentrations of structural Cr in olivine may be beyond the
detection limit of the electron microprobe. Most significantly, there is ample room in the various cavities
combined for a large quantity of volatiles.
Roedder (1965, 1984) identified two types of inclusions in olivine: `primary', which were incorporated
directly into growing crystal faces, and `secondary',
which were introduced into the mineral grains along
straight or curvilinear fractures that are now annealed.
The latter are clearly important in the Juan Fernandez
olivine. However, the olivine also contains faceted
cavernous inclusions, ellipsoidal inclusions, rounded
multiphase inclusions, tiny single-phase or multiphase
inclusions with preferential crystallographic orientation, cavities adjacent to spinel, and tiny faceted pits.
These are all forms of primary inclusions. They are
so diverse that this single term does not do them
descriptive justice, although it does connote the
common aspect of their origin.
Drever & Johnston (1957) emphasized mechanisms
of crystal growth, rather than resorption, as the most
important factor in producing irregular morphologies
among olivine phenocrysts. Experimental studies
(Donaldson, 1976) extended this conception to the
scale of tiny olivine with elongate, strikingly faceted
or highly decorated morphologies that grew at exceedingly high undercooling, as for example in lunar samples or the margins of pillow basalts. The experiments
showed that undercooling is systematically related to
the diverse morphologies olivine exhibits from the
interior of flows to their quenched margins.
The same range of faceted morphologies exists
here, but in reciprocal form, in the various primary
cavities that surround inclusions, whether they are
empty or filled with glass and minerals, among the
Juan Fernandez olivine crystals. The tiniest ones are
like the most elaborate hopper olivine in pillow margins. Somewhat larger ones have serrated boundaries
of the type typical of skeletal or coarsely dendritic
overgrowths. The largest ones may either be faceted
or they are rounded, as are the large re-entrants in the
coarsely skeletal olivine phenocrysts illustrated by
Drever & Johnston (1957), and the similar and even
larger `harrisitic' olivines, which are coarsely skeletal
crystals orthogonal to layering in gabbroic intrusions,
as discussed by Donaldson (1982). The only real distinction between skeletal re-entrants and the large cavities is that in most cases the cavities are empty or only
partly filled. I consequently believe that all the primary inclusions described here have morphologies
indicating crystal growth around them at a range of
undercooling.
Despite this, as all cavities occur in otherwise euhedral or tabular olivine, the general characteristics of
crystal growth had to be at small undercooling, at rates
of perhaps 02±04 mm/s (Donaldson, 1975), which
would produce a typical euhedral phenocryst in a few
days. Thus each re-entrant or faceted cavity represents
a local spot on the crystal that experienced heightened
undercooling.
The key to this is the presence of bubbles. Theory
predicts that when the surface energy between vapor
and crystal is less than that between vapor and melt,
then the activation energy for nucleation on the crystal
surface is less than in the melt, thus the crystal surface
is where bubbles will preferentially attach (Sigbee,
1969). Bubble nucleation consequently is dominantly
heterogeneous, and favored by both the presence of
crystals in the melt where bubbles attach, and the
roughness of the crystal surfaces (Bagdassarov &
Dingwell, 1993). The roughness reduces the activation
energy needed for nucleation even further than the
439
JOURNAL OF PETROLOGY
VOLUME 44
simple presence of crystals. Observations on silicic
melts also indicate that bubbles may also nucleate
preferentially on certain minerals, in particular the
Fe±Ti oxides (Hurwitz & Navon, 1994; Navon &
Lyakhovsky, 1998). In Juan Fernandez tholeiitic
picrites, another oxide mineral, namely spinel included
in olivine, is almost invariably associated with a
bubble, but whether this has to do with the spinel
adding roughness to the surface of the growing olivine
or a predilection for bubbles to nucleate on the spinel is
difficult to say.
Whether or not spinel is present, bubbles that nucleate on olivine surfaces would in the first instance be
local sites without molten silicate to supply the growing
olivine. The olivine would tend to grow around them.
In the second place, the bubbles represent places where
local liquidus conditions are modified. Lipman et al.
(1985) and Lipman & Banks (1987) argued that
increases in microphenocryst content during the eruption at Mauna Loa in 1984, but without a change in
lava temperature, resulted from undercooling of
20±30 below the liquidus as a result of vesiculation
and volatile release that occurred as the lava flowed
down the mountain (see Cashman & Mangan, 1994).
Release of volatiles simply elevates liquidus boundaries, driving melts from conditions that produce tabular crystals to those that produce skeletal or dendritic
crystals. On the surfaces of growing phenocrysts,
bubble nucleation would also modify melt viscosities,
thus changing diffusion gradients in the layer of melt
immediately around the crystals.
Growing crystals also produce their own supercooling effect when there is material in the melt that is
more soluble than in the crystal. This is called constitutional supercooling, and it results in the concentration of the impurity being greatest near the crystal (e.g.
Kirkpatrick, 1975). Both the liquidus temperature and
the undercooling increase away from the crystal; with the
result that protuberances on the crystal grow faster
the further they are from the main crystal interface.
The tendency to faceted morphologies is accentuated.
Regular spacing of inclusions and cavities is important. Kirkpatrick (1975, 1981) emphasized that there is
as yet no adequate theory to account for planar interface stability during crystal growth, despite the fact
that euhedral crystals are common. Existing theory
predicts that, with increasing undercooling, interface
instability should set in at some point controlled by the
crystal growth rate Y and the rate of diffusion D of the
rate-controlling component of the melt (Cahn, 1967).
When D/Y is small, large euhedral crystals can form.
As D/Y increases, planar interfaces are no longer
stable. The first instabilities to appear have long wavelengths, resulting in faceted crystals. As D/Y increases
still further, the spacing of instabilities decreases;
NUMBER 3
MARCH 2003
morphologies are dendritic or skeletal. At very high
undercooling, and greatest D/Y, morphologies are
branching (spherulitic). Relationships between undercooling and morphologies are borne out by a
number of petrographic and experimental studies
(e.g. Kirkpatrick, 1979; Lofgren, 1983). Implicitly,
the spacing of instabilities, and thence of protuberances
on planar interfaces, is regular at a given undercooling
(Keith & Padden, 1963).
Thus the regular spacing of dendrites on olivine
phenocrysts (Fig. 5e and f ), of ellipsoidal cavities
(Fig. 7a), of inclusion dendrite spacings (Fig. 7h),
and of rows of tiny faceted cavities (Fig. 11) are
probably aspects of development of planar instabilities
during the heightened undercooling occasioned by
sudden cooling or vesiculation. In the case of tiny
faceted cavities, the instabilities developed at a characteristic spacing of 3±5 mm, and determined where
tiny bubbles nucleated. Given the otherwise euhedral
morphology of the olivine enclosing these cavities, this
may represent only incipient breakdown of interface
stability, at small undercooling that would not be
apparent at all except for the bubbles that formed. If
this is the case, then the spacing may represent the
critical distance for the onset of interface instability of
olivine in these materials (Cahn, 1967).
The presence of non-uniform boundary layers
around the growing crystals is also important. At
somewhat elevated undercooling promoted by vesiculation, spots near where bubbles are attached to an
olivine would not experience as efficient an exchange
of components into and away from the growing crystal
as when nearer equilibrium; constitutional supercooling would not be uniform. Conversely, because of very
local viscosity contrasts in the boundary layer, such
exchange would be more efficient on portions of crystal
surfaces away from bubbles. The surface energy of the
interface varies, and this controls the shape of surface
instabilities (Cahn, 1967). The bubbles therefore
inhibit diffusion parallel to the crystal interface, this
tending to promote crystal growth instabilities.
In general, melt species that are particularly incompatible with the olivine structure should tend to build
up irregularly in the boundary layer, being more concentrated nearer the bubbles. These would include
alkalis, S, Cr, and volatiles, and would accentuate
any tendency for irregularities on the olivine surfaces
to act as sites for spinel and sulfide nucleation. This
may account for the unusual concentration of Cr-spinel
and, in two samples, sulfides, even in large multiphase
inclusions.
In crystalline basalt, there is no way to determine
whether such boundary layers existed. Glass next to
quenched olivine in pillow rims might have a somewhat different composition than glass further away,
440
NATLAND
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
Fig. 13. Tiny sulfides in boundary layers near olivine quench crystals and phenocrysts in the glassy margin of sample SD-7C, a picritic
tholeiite from Siqueiros Fracture Zone, eastern Pacific (Natland, 1989). The photomicrographs are in oil-immersion reflected light. The
olivine in (a) is a typical hopper crystal; that in (b) is more equant, but it has small dendritic prongs. The tiny bright specks are minute sulfides,
concentrated in boundary layers adjacent to the olivines. The sulfides in (c) are in clear glass (black) between a spinel (light gray) and an
olivine (dark gray). The spinel, sulfides, and glass are in a skeletal embayment in a large olivine phenocryst. In (d), the tiny sulfides are
concentrated near the straight edge of a large, faceted, olivine phenocryst.
but no one has as yet attempted to prove the existence
of boundary layers in natural glasses using an electron
microprobe, although some of the potential inhomogeneities perhaps have shown up in analyses of glass
inclusions. For example, Dixon et al. (1991) reported S
contents among Puna Ridge glass inclusions in olivines,
other minerals, and gabbro xenoliths that are in excess
of concentrations necessary for sulfide saturation.
These occur even in basalts with quenched glasses not
saturated in sulfide. Supersaturation in S can also be
inferred using data from inclusions of the 1959 Kilauea
summit eruption (Anderson & Brown, 1993). Dixon
et al. (1991) speculated that the excess results from
post-entrapment growth of the surrounding minerals.
However, they did not consider the possibility of
boundary-layer buildup during skeletal crystal
growth, and did not report whether there are any
very tiny sulfide globules in their inclusions like
those in Fig. 8.
Degassed basalts erupted subaerially or in shallow
water do not have such high concentrations of sulfur,
and thus are not likely to provide evidence for a boundary layer enriched in S or sulfide next to a crystal
surface of a quenched olivine. Indeed, in most Juan
Fernandez tholeiites, sulfide is not present even in
inclusions in olivine phenocrysts, suggesting that S
was almost completely lost to degassing before their
crystallization. Nevertheless, as sulfide is present in
inclusions of two samples, I have examined quenched
glass next to olivine in the pillow rim of a primitive,
sulfide-saturated picritic abyssal tholeiite dredged from
Siqueiros Fracture Zone, eastern Pacific, using highmagnification (1600) oil-immersion reflected light.
The specimen has been described in more detail elsewhere (Natland, 1980, 1989). This glass, indeed, has
myriads of tiny sulfide globules in boundary layers next
to hopper olivine (Fig. 13a and b) and faceted olivine
phenocrysts (Fig. 13c and d). Most of the sulfides in
441
JOURNAL OF PETROLOGY
VOLUME 44
these photomicrographs show up because of internal
reflection, and the grains are so tiny that they produce
a noticeable asterism at this high magnification. Nevertheless, they are clearly much more reflective than
either the olivine or the quench clinopyroxene dendrites in Fig. 13a and in transmitted light they appear
as tiny black specks. The sulfide in Fig. 13c is in clear
glass between an olivine phenocryst and a spinel grain
contained within a skeletal embayment in the olivine.
This confirms that boundary-layer segregation of sulfide can occur at the high undercooling characteristic
of pillow margins near growing olivine, and thus need
not result from closed-system post-entrapment crystallization of host olivine, although it does not prove
that it occurred in sulfide-saturated basalts from either
Juan Fernandez or Kilauea. Heightened undercooling
resulting from bubble formation also explains why many
melt inclusions retain glass, or have clinopyroxenes
with elaborate dendritic morphologies, even in lavas
with a holocrystalline groundmass. Such morphologies
are at odds with Roedder's (1965) inference that
bubbles within melt inclusions were not there originally, but that they nucleated and grew as the inclusions
slowly crystallized. Instead, the bubbles, like the glass
and its elaborate silicate minerals, were quenched in
place. Occurrence in a thin section of a population
of olivine phenocrysts with variable proportions of
inclusions is a reflection of uneven distribution of
volatiles exsolving from a melt in a magma body at
large (e.g. Kennedy, 1955; see also Farley &
Craig, 1994, for discussion pertaining to Ar concentrations), and from processes acting at crystallization
interfaces.
Indeed, crystal surfaces of olivine may not have been
the exclusive sites of bubble nucleation. Bubbles may
instead have nucleated in the melt, perhaps in response
to magma stirring or mixing, and simply attached to
olivine crystals as they streamed past. This would have
produced uneven bubble distribution on the olivine,
and variable concentrations of inclusions in different
olivine crystals of the same composition, now present in
a single lava sample. Large bubbles would not have
remained attached, their buoyancy being too great
for the effects of surface tension at the bubble±crystal
interface to overcome. Large discontinuities in mineral
surfaces, such as intersections of crystal faces, would
probably have aggregated larger bubbles from smaller
ones moving across the crystal surfaces, and been principal loci of escape of buoyant bubbles (Fig. 12a). The
attached bubbles might have been sufficient to overcome the density contrast between olivine and liquid,
and thus to retard or eliminate the tendency for olivine
to sink. Attached (as opposed to directly nucleated)
bubbles would also have physically penetrated any
boundary-layer melt, perturbing not only the growth
NUMBER 3
MARCH 2003
of the crystal, but the chemistry of the boundary layer
as well.
Inclusions along annealed secondary curvilinear
fracture surfaces contain the same phases as primary
multiphase inclusions, although particular phasesÐ
spinel, sulfide, and volatilesÐtend to be concentrated
along different portions of them. As mentioned earlier,
this suggests a mechanism of physical separation of
these phases during flow along the microfractures.
Sulfides and volatiles appear to have been exceptionally mobile and easy to separate from silicate melt, the
former in the form of an immiscible liquid, the latter as
either a two-phase or a supercritical fluid. The unusual
concentrations of spinel, sulfide, and volatiles (in the
form of empty cavities) along these healed microfractures point to incorporation of the same boundary
layer material that was entrapped during normal
crystal growth by larger multiphase inclusions. The
fractures formed at the same time that vesiculation
introduced local perturbations in the properties of
the boundary layer melts. Differential stresses on
the mineral surface resulting from contact with melt
at different undercooling, or from contrasting effects
of surface tension at bubble±melt contacts, may
have produced some of the fractures. However, most
of the fractures are probably simple contraction
features resulting from sudden cooling of the growing
olivine phenocrysts during magma mixing. Additional
indications of mixing and some of the conditions of
vesiculation are provided by mineral compositions,
considered next.
MINERAL COMPOSITIONS
Tables 1 and 2 give compositions of olivine and spinel
analyzed in eight Juan Fernandez olivine tholeiites.
Olivine cores and rims together range in composition
from Fo88 to Fo58, and have an asymmetric distribution about a single prominent mode at Fo82 (Fig. 14a).
All values are sufficiently iron rich for the olivine to
be phenocrystic. Spinel ranges from magnesiochromite
to titanian chromite in composition. It is strongly
bimodal in composition (Fig. 14b), and its Mg-number
correlates with the Fo content of host olivine (Fig. 14c).
The bimodality in spinel defines two sample groups
intergrown with olivine, Group 1 with more, and
Group 2 with less, magnesian olivine and spinel
(Fig. 14a±c). Some even more iron-rich spinel is not
intergrown with olivine. Within each rock, there is a
range of mineral compositions, but no sample has both
spinel and olivine bridging between Groups 1 and 2.
Most spinel is more chromian, but less magnesian,
and all is more titanian, than spinel in abyssal tholeiites
(Fig. 14d and e), in abyssal and alpine peridotites (e.g.
442
NATLAND
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
Table 1: Electron microprobe analyses of olivine and their estimated equilibrium liquid compositions and
temperatures
Sample
Run-analysis*
Typey
Fo
MgL
T ( C)KILz
99.88
98.88
81.3
80.6
44.9
43.3
1147
0.27
0.27
97.25
97.51
81.5
81.5
45.3
45.3
1148
0.20
0.19
0.26
0.12
100.69
98.90
80.0
66.3
42.0
20.7
1142
41.81
30.86
0.20
0.26
0.28
0.08
98.96
99.71
80.1
63.4
42.2
17.5
1142
0.60
0.52
32.39
33.82
0.26
0.22
0.16
0.23
100.51
100.47
65.4
67.8
19.7
22.4
1092
35.51
20.02
0.60
0.24
28.3
41.04
0.24
0.20
0.14
0.27
100.34
100.00
58.7
78.5
13.2
38.9
1075
39.09
39.91
17.69
17.38
0.27
0.24
43.07
42.95
0.22
0.22
0.28
0.23
100.62
100.33
81.3
81.5
44.9
45.3
1147
38.69
38.47
19.16
19.75
0.28
0.22
41.66
40.79
0.18
0.19
0.25
0.27
100.22
99.69
79.5
78.5
41.0
39.1
1140
39.17
38.76
16.19
16.95
0.25
0.25
43.00
43.23
0.21
0.25
0.28
0.23
99.82
99.67
81.9
82.0
46.3
46.5
1150
37.06
37.55
24.79
23.79
0.37
0.36
37.02
36.39
0.24
0.20
0.13
0.16
99.61
98.45
72.7
73.2
29.0
29.7
1114
17.30
24.96
0.22
0.37
43.40
37.66
0.23
0.20
0.29
0.14
100.90
101.00
81.7
72.9
45.8
29.3
1149
SiO2
FeO
MnO
MgO
CaO
NiO
38.65
38.05
17.57
18.11
0.26
0.24
42.96
42.08
0.18
0.18
0.26
0.22
37.21
37.23
17.13
17.17
0.23
0.23
42.22
42.38
0.19
0.23
38.85
35.90
18.86
29.60
0.26
0.43
42.26
32.66
37.96
36.16
18.51
31.76
0.20
0.59
36.61
36.99
30.49
28.69
35.55
38.23
Total
Alexander Selkirk (Mas Afuera)
MF C-2
2-05
PH 1
MF C-2
2-08
PH 2
MF C-2
2-10
PH 3
MF C-2
2-11
PH 4
MF C-2
2-13
PH 5
MF C-2
2-19
PH 6R
MF C-2
2-20
PH 6C
MF C-2
6-20
GM
MF C-2
6-21
R
MF C-2
6-22
R
MF C-2
6-25
GM
MF C-4
4-25
PH 1-1
MF C-4
4-27
PH 1-2
MF C-4
4-29
PH 1-3
MF C-4
4-31
PH 1-4
MF C-4
4-33
PH 1-5
MF C-4
4-35
PH 1-6
MF C-4
1-38
PH 2
MF C-4
1-38
SK 1C
MF C-4
1-39
SK 1R
MF C-4
1-43
PH 3
MF C-4
1-45
PH 4
39.46
37.67
MF C-4
1-46
PH 5
36.88
17.24
0.28
42.41
0.20
0.27
99.28
81.4
45.1
1148
16.32
21.14
0.30
0.29
43.58
39.69
0.13
0.28
0.36
99.71
99.88
82.6
77.0
48.0
1153
1130
20.48
17.49
0.31
0.32
40.03
42.87
0.29
0.27
99.74
100.28
77.7
81.4
36.1
37.4
1148
19.74
18.66
0.32
0.42
41.32
41.22
0.24
0.20
100.31
99.11
78.9
79.7
45.1
39.7
1140
26.72
36.32
0.45
0.55
35.50
27.00
0.31
0.44
100.37
99.97
70.3
57.0
41.4
25.6
1072
16.30
16.80
0.21
0.33
42.93
43.02
0.25
0.09
98.72
99.33
82.4
82.0
11.7
47.5
1150
15.66
17.64
0.22
0.23
44.16
42.52
0.27
0.27
99.31
99.28
83.4
81.1
46.5
50.0
1146
17.06
15.23
0.27
0.23
43.32
43.72
0.28
0.25
99.86
98.23
81.9
83.7
44.4
46.3
1170
15.77
15.60
0.20
0.19
44.34
43.63
0.38
0.26
99.70
98.01
83.3
83.3
50.8
49.7
1156
17.77
19.06
0.27
0.34
43.13
40.39
0.28
0.33
100.35
98.56
81.2
79.1
49.7
44.7
1138
20.45
17.59
0.39
0.30
40.10
42.24
0.36
0.31
77.8
81.1
0.37
40.82
0.32
99.26
79.1
44.4
40.1
1146
19.21
0.19
0.13
99.74
99.31
40.1
37.6
14.40
0.25
45.51
0.36
0.28
100.46
84.9
54.1
1140
MF 20
6-03
PH 1
39.02
MF 20
6-06
D
MF 20
6-07
w pl
38.26
38.44
MF 20
6-09
PH 2
MF 20
6-11
SK
MF 20
6-12
PH 3C
MF 20
6-13
PH 3R
MF 20
6-15
GM(eu)
MF 20
6-16
PH 4
MF 20
6-19
PH 5
MF 3
2-21
PH 1
MF 3
2-22
GM
MF 3
2-23
PH 2
MF 3
2-25
PH 3
MF 3
2-46
PH 4
MF 3
2-48
PH 5
MF 3
6-29
GM
MF 3
6-30
GM
MF 3
6-31
GM
MF 3
6-32
GM
MF 3
6-33
GM
38.68
38.41
MF S-1
1-04
M1 G1
39.76
39.13
38.48
38.35
37.22
35.59
38.76
38.85
38.74
38.39
38.73
38.51
38.84
38.06
38.63
38.30
38.26
443
0.22
0.19
0.20
0.21
0.26
0.17
0.07
0.27
0.24
0.26
0.23
0.20
0.29
0.27
0.25
0.22
0.14
0.16
1144
1148
1095
1087
1099
1136
1148
1136
1150
1116
1115
1132
1137
1107
1152
1164
1150
1156
1147
1133
1138
JOURNAL OF PETROLOGY
VOLUME 44
NUMBER 3
MARCH 2003
Table 1: Continued
Fo
MgL
T ( C)KILz
99.53
99.24
85.0
86.90
54.3
60.1
1141
0.30
0.28
99.13
98.47
86.5
86.9
58.8
60.1
1153
0.31
0.22
0.23
0.34
100.37
99.69
83.7
88.2
50.8
64.5
1132
45.6
46.08
0.24
0.30
0.26
0.32
99.73
100.37
85.7
85.8
56.4
56.7
1147
0.24
0.26
45.97
44.29
0.26
0.25
0.22
0.21
99.85
100.08
85.2
83.4
54.9
50.0
1143
14.51
16.52
0.24
0.21
45.03
43.96
0.26
0.32
0.19
0.21
99.72
100.65
84.7
82.6
53.5
48.0
1139
37.66
38.81
21.77
16.86
0.29
0.25
38.32
42.83
0.46
0.35
0.19
0.22
98.69
99.32
75.8
81.9
34.0
46.3
1066
39.49
38.74
15.19
20.76
0.22
0.25
44.70
40.58
0.28
0.38
0.29
0.18
100.17
100.89
84.0
77.7
51.6
37.4
1134
39.28
39.5
14.99
16.17
0.19
0.21
44.74
44.34
0.28
0.23
0.30
0.36
99.78
100.81
84.2
83.0
52.1
49.0
1135
39.11
38.787
16.39
16.66
0.22
0.27
43.28
43.88
0.31
0.25
0.26
0.28
99.57
100.12
82.5
82.4
47.7
47.5
1123
38.79
39.21
17.75
14.8
0.18
0.21
41.96
45.07
0.33
0.27
0.15
0.26
99.16
99.82
80.8
84.4
43.8
52.7
1113
39.03
38.74
18.19
17.31
0.27
0.24
40.72
42.22
0.52
0.32
0.19
0.26
98.62
99.09
79.9
81.3
41.8
44.9
1107
38.41
38.74
16.41
16.65
0.21
0.32
42.32
42.35
0.39
0.39
0.21
0.33
97.95
96.78
82.1
81.9
46.8
46.3
1121
38.45
38.79
16.74
15.19
0.27
0.21
42.53
44.52
0.36
0.30
0.21
0.29
98.56
99.30
81.9
83.9
46.3
51.3
1119
39.00
38.44
12.37
13.07
0.20
0.26
45.89
45.11
0.24
0.33
0.34
0.27
98.04
97.48
86.9
86.0
60.1
57.3
1164
38.57
38.12
15.17
13.82
0.24
0.20
44.40
44.59
0.31
0.27
0.27
0.27
98.96
97.27
83.90
85.2
51.30
54.9
1133
39.14
38.75
17.57
14.07
0.35
0.25
42.37
45.05
0.40
0.29
0.26
0.23
100.09
98.64
81.1
85.1
44.4
54.6
1114
38.50
38.39
15.52
16.89
0.24
0.28
43.57
42.65
0.34
0.36
0.21
0.22
98.38
98.79
83.3
81.8
49.7
46.0
1129
37.97
38.00
19.22
17.02
0.27
0.28
40.54
42.41
0.39
0.38
0.19
0.18
98.58
98.27
79.0
81.6
39.9
45.6
1102
37.79
38.64
12.72
15.77
0.19
0.36
45.09
43.23
0.26
0.35
0.30
0.16
96.35
98.51
86.3
83.0
58.2
49.0
1151
39.08
39.02
16.79
15.58
0.36
0.27
43.02
43.46
0.34
0.36
0.26
0.22
99.85
98.91
82.1
83.2
46.8
49.5
1121
Sample
Run-analysis*
Typey
SiO2
FeO
MnO
MgO
CaO
NiO
MF S-1
1-08
M1 G2
MF S-1
1-09
M2 G1
39.01
39.45
14.24
12.52
0.20
0.23
45.37
46.46
0.40
0.28
0.31
0.30
MF S-1
1-11
M2 G2-1
MF S-1
1-12
M2 G2-2
39.28
39.39
12.64
12.35
0.19
0.19
45.32
46.02
0.40
0.24
MF S-1
1-13
M3 G3
MF S-1
1-14
M3 G2
39.33
40.05
15.52
11.35
0.23
0.17
44.75
47.56
MF S-1
1-15
M4 G1
MF S-1
1-17
M4 G2
39.90
39.96
13.55
13.53
0.18
0.18
MF S-1
1-18
PH 1C
MF S-1
1-19
PH 1R
38.96
39.36
14.20
15.71
MF S-1
1-21
2
MF S-1
1-22
GM 1
39.49
39.41
MF S-1
1-23
GM 2
MF S-1
1-24
PH 3
MF S-1
1-25
PH 4
MF S-1
1-26
GM 4
MF S-1
1-27
PH 5C
MF S-1
1-29
PH 6
MF S-1
1-30
PH 7
MF S-1
1-31
H8
MF S-1
1-32
GM 5
MF S-1
1-33
PH 9
MF S-1
7-25
GM 3
MF S-1
7-27
GM 4R
Total
1164
1164
1213
1147
1129
1124
1119
1096
1127
1123
1137
1116
Robinson Crusoe (Mas a Tierra)
PIN 5
6-38
GM 1
PIN 5
6-39
GM 2
PIN 5
6-40
GM 3
PIN 12
2-27
MPH 1
PIN 12
2-30
MPH 2
PIN 12
2-32
MPH 3
PIN 12
2-34
MPH 4
PIN 12
2-35
MPH 5
PIN 12
2-36
GM 2R
PIN 12
2-38
MPH 6
PIN 12
2-40
MPH 7
PIN 12
2-42
PH 1-1
PIN 12
2-43
PH 1-2
PIN 12
2-44
PH 2
PIN 12
2-45
PH 3
PIN 12
6-34
GM 1
PIN 12
6-35
GM 2
PIN 12
6-37
GM 3
1119
1133
1149
1143
1142
1119
1118
1127
1128
*Run-analysis: electron microprobe runs on different dates and analysis number on that date; run given to distinguish
analyses with the same number.
yType: PH, phenocryst; MPH, microphenocryst; GM, groundmass; C, core; R, rim; M, grain mount; G, grain; Ð, spot; E,
euhedral; D, dendrite; SK, skeletal grain. Number after PH, MPH, GM or M identifies different grains.
zT ( C)KIL is crystallization temperature calculated using modified Kilauea geothermometer as described in the text.
444
Table 2: Electron microprobe analyses of spinel, compositional parameters, and estimated equilibrium liquid compositions and temperatures compared
with those of adjacent intergrown olivine and its Fo content
Sample
Run-analysis*
TiO2
4.73
4.45
2-07
MF C2
2-09 PH2
MF C2
2-14 PH5
MF 3
2-26
MF 3
2-49
MF C-4
4-24 PH1-1
MF C-4
4-26 PH1-2
MF C-4
4-28 PH1-3
MF C-4
4-30 PH1-4
MF C-4
4-32 PH1-5
MF C-4
4-34 PH1
MF C-4
4-35 PH1
MF C-4
4-36 PH1
MF C-4
1-36 GM
MF C-4
4-37 PH2
MF C-4
1-39
MF C-4
1-40 SK1R
MF C-4
1-41 GM
MF C-4
1-42 PH3
MF C-4
1-45 PH4
MF C-4
1-47 PH5
MF C-4
1-48 SK2
MF 20
6-04 PH1
MF 20
6-05 PH1
MF 20
6-10 PH2
MF 20
6-16-4
MF 20
6-17 PH4
MF S-1
1-05 M1G1
MF S-1
1-07 M1G1
MF S-1
1-10 M2G1
2.33
2.40
MF S-1
1-16 M4G1
2.92
7.22
3.64
2.77
5.63
3.79
3.63
5.61
5.65
7.94
5.55
3.62
4.56
6.62
9.07
7.86
14.14
5.75
7.86
7.16
8.27
5.44
6.65
6.20
6.16
4.69
2.24
Cr-no.z
Fe-no.z
Fo
MgL(ol)x
T ( C)Olx
MgL(Sp){
T ( C)Sp{
0.25
0.23
100.37
100.81
39.0
41.1
54.3
57.3
17.9
17.7
81.3
56.6
1147
45.7
48.3
1118
9.76
8.38
0.27
0.28
100.57
99.89
42.2
35.4
46.4
49.1
15.9
21.9
80.6
80.0
55.3
54.5
1144
45.4
42.5
1117
0.30
0.21
11.62
111.66
0.24
0.28
99.23
98.39
51.3
52.8
54.4
54.1
12.4
11.1
83.6
83.3
60.5
59.9
1168
55.6
56.0
1144
37.67
34.33
0.28
0.33
9.58
9.84
0.30
0.20
100.63
102.15
40.6
42.9
43.4
51.8
19.0
81.2
78.5
15.5
52.3
56.6
1136
44.4
47.2
1114
23.54
26.43
34.20
38.48
0.28
0.28
10.08
9.04
0.23
0.26
101.80
101.32
43.5
38.0
44.2
41.9
15.0
18.2
81.5
79.5
56.9
53.8
1148
45.1
41.2
1116
13.49
16.05
25.66
28.83
37.79
43.07
0.26
0.36
9.09
8.00
0.26
0.26
99.36
100.24
38.9
33.4
42.0
49.6
18.8
24.6
78.6
52.4
1136
42.3
41.6
1109
23.96
23.28
15.10
11.96
25.99
22.98
39.57
33.75
0.34
0.32
8.79
9.55
0.28
0.27
100.27
97.24
37.8
42.7
44.2
38.2
21.0
15.7
42.4
42.8
1109
18.78
18.74
29.01
24.52
12.67
14.54
23.06
26.13
34.45
39.21
0.32
0.26
9.66
9.26
0.14
0.24
98.20
100.31
42.9
38.9
50.9
46.8
17.5
20.9
48.1
44.9
1124
11.69
12.03
19.89
22.30
20.54
19.53
34.10
32.80
52.58
50.17
0.32
0.40
4.78
4.88
0.20
0.19
100.59
99.79
20.1
21.1
53.3
55.4
34.4
31.6
29.6
30.4
1075
2.19
19.69
2.34
27.64
36.62
11.24
40.08
25.50
73.02
35.61
0.33
0.36
2.22
9.12
0.16
0.22
98.06
99.72
9.1
39.1
41.7
48.7
86.1
15.7
21.6
43.6
1053
13.61
15.13
23.69
20.91
16.57
17.24
31.84
29.34
46.74
44.85
0.35
0.28
5.55
6.31
0.25
0.25
99.72
96.62
23.8
27.9
53.9
48.1
26.4
27.4
1081
11.83
11.60
19.92
31.30
32.07
18.48
49.34
23.80
0.33
40.42
4.84
0.33
0.17
9.29
96.63
0.27
21.3
100.51
53.0
41.2
32.7
64.4
11.73
18.99
28.97
27.65
18.08
11.40
24.52
24.22
40.78
34.47
0.37
0.30
9.37
9.96
0.30
0.34
99.89
99.06
40.7
42.5
62.3
49.4
16.44
18.33
27.36
28.27
14.19
14.36
28.12
22.83
40.89
35.75
0.46
0..37
7.35
9.90
0.21
0.21
100.29
98.96
31.9
43.8
52.8
50.9
18.20
19.19
41.52
40.83
7.28
6.63
17.73
17.75
24.28
23.71
0.25
0.26
12.00
12.12
0.24
0.25
99.46
99.36
54.8
55.1
60.5
56.8
19.94
18.82
41.53
40.96
5.82
5.72
17.25
17.52
22.49
22.67
0.21
0.29
12.81
12.54
0.16
0.12
100.12
98.89
57.1
56.2
58.3
59.4
Fe2O3y
FeOy
FeOTy
MnO
MgO
17.05
16.37
30.19
32.80
12.70
12.97
25.25
24.07
36.68
35.74
0.35
0.27
9.01
9.37
21.74
17.02
28.08
24.48
12.04
14.67
23.98
27.47
34.81
40.67
0.25
0.37
19.44
20.35
34.55
35.76
9.44
8.65
19.80
18.71
28.29
26.49
21.37
20.15
24.41
32.30
13.66
12.04
25.20
23.50
23.98
22.33
28.31
23.97
11.85
13.40
21.63
15.82
23.32
23.18
20.26
25.26
NiO
81.9
82.0
57.6
57.7
1142
1156
1147
1140
1150
1150
1125
1109
1145
1122
1106
1107
1110
1116
1077
81.7
72.9
57.3
44.7
1149
81.4
56.6
1148
32.2
35.6
30.6
53.0
1077
1136
1115
1105
1091
26.6
27.1
82.6
82.6
58.7
58.7
1153
16.2
20.7
81.4
56.8
1148
52.6
47.7
1099
19.7
9.2
82.4
84.9
58.4
62.8
1152
38.8
49.6
1154
8.3
7.2
84.9
86.5
62.8
66.0
1194
59.3
58.8
1169
7.3
85.7
64.3
1210
60.5
60.1
1153
1194
1229
1138
1123
1128
1153
1167
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
445
MF C2
Mg-no.z
Cr2O3
NATLAND
Alexander Selkirk (Mas Afuera)
MF C2
2-06 PH1
5.57
Total
Al2O3
Table 2: Continued
Run-analysis*
MF S-1
1-20 PH1C
MF S-1
1-22 PH2
MF S-1
1-24 PH1
MF S-1
1-26 PH4
MF S-1
1-28 PH5
MF S-1
1-34 PH9
TiO2
Mg-no.z
Cr-no.z
Fe-no.z
Fo
MgL(ol)x
T ( C)Olx
MgL(Sp){
T ( C)Sp{
0.17
0.26
100.39
101.34
54.8
51.9
53.9
64.1
6.5
11.0
85.2
84.7
63.3
62.4
1200
55.9
58.1
1145
12.82
11.68
0.24
0.18
100.04
98.50
57.3
53.3
52.3
58.0
7.4
9.5
85.2
84.0
63.3
61.2
1200
58.1
57.2
1151
0.27
0.32
11.83
11.74
0.24
0.18
99.83
98.92
53.1
53.0
56.1
53.9
9.9
8.7
84.2
84.5
61.5
62.1
1180
56.5
55.2
1147
27.87
27.43
0.23
0.26
11.29
11.24
0.19
0.24
100.93
100.99
50.7
50.5
61.4
62.5
11.6
10.9
83.2
83.7
59.8
60.6
1158
56.5
56.6
1140
17.37
15.07
23.52
18.12
0.26
0.21
12.33
12.77
0.19
0.20
101.01
95.40
56.0
60.3
61.7
61.2
8.4
4.4
86.0
86.0
64.8
64.8
1216
60.4
83.4
1168
7.17
3.93
18.73
16.52
25.18
20.05
0.28
0.25
11.46
12.58
0.18
0.17
99.15
98.70
52.3
57.8
59
9.1
4.9
83.9
86.9
61.0
67.0
1174
56.4
61.2
1147
43.22
40.26
6.12
7.83
19.25
12.88
24.76
25.26
0.26
0.26
11.14
11.55
0.21
0.18
99.76
98.72
50.9
53.2
7.8
10.0
86.0
84.1
64.8
61.3
1216
56.1
57.1
1146
42.08
39.60
7.09
7.87
17.06
19.52
23.44
26.60
0.26
0.29
12.14
11.20
0.21
0.18
99.03
99.84
56.1
50.7
8.9
10.0
85.1
83.4
63.1
60.1
1198
60.3
55.2
1166
Fe2O3y
FeOy
FeOTy
MnO
1.71
2.09
22.82
16.30
39.77
43.41
5.39
8.81
18.05
18.85
22.90
26.78
0.28
0.28
12.20
11.34
1.79
2.46
23.45
19.03
38.25
39.15
6.14
7.44
17.14
18.36
22.66
25.05
0.21
0.220
2.67
2.53
20.02
21.33
38.16
37.19
7.88
6.92
18.76
18.71
25.84
24.93
17.10
16.65
40.50
41.36
9.07
8.53
19.71
19.76
18.30
18.50
43.94
43.42
8.85
3.39
18.70
19.38
40.22
44.06
17.04
18.18
18.36
18.52
Robinson Crusoe (Mas a Tierra)
6-24
2.84
2.95
1.77
PF 5
6-26
PF 5
6-28
PF 5
6-30
PIN 12
2-28 MPH1
PIN 12
2-31 MPH2
PIN 12
2-33 MPH3
PIN 12
2-37 MPH
PIN 12
2-39 MPH6
42.24
1.83
PIN 12
2-41 MPH7
2.66
1.84
2.41
1.81
2.52
MgO
NiO
60.40
63.0
59.8
60.6
58.9
1190
1175
1186
1170
1216
1240
1177
1164
1151
1149
1143
1147
1201
1176
1148
1143
NUMBER 3
*Run-analysis: electron microprobe runs on different dates and analysis number on that date; run given to distinguish analyses with the same number. Remaining notation
identifies intergrown olivine phenocryst and microphenocryst as in the second footnote to Table 1.
yFe2O3 and FeO are calculated from total iron as FeOT by stoichiometry.
zMg-number ˆ Mg/(Mg ‡ Fe2‡); Cr-number ˆ Cr/(Cr ‡ Al); Fe-number ˆ Fe3‡/(Fe3‡ ‡ Cr ‡ Al) from structural formulae.
xMgL(ol) is calculated using KD values as discussed in the text. T ( C)Ol is calculated using MgL and the modified Kilauea geothermometer of Helz & Thornber (1987) as
discussed in the text.
{MgL(Sp) and T ( C)Sp are calculated using the procedure of Allan et al. (1988, 1989).
VOLUME 44
Cr2O3
PF 5
446
Total
Al2O3
JOURNAL OF PETROLOGY
Sample
MARCH 2003
NATLAND
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
Dick & Bullen, 1984), and in residual Pacific mantle
xenoliths (e.g. Dieu, 1995). This and the phenocrystic
olivine is sufficient to support the initial thesis that
mantle helium was trapped in phenocrysts in the
cavities described, rather than in minerals extracted
from mantle wall rocks. On the other hand, the Juan
Fernandez spinel is similar to spinel associated with
olivine phenocrysts in Puna Ridge tholeiites (Clague
et al., (1995), although the most magnesian is not quite
as chromian. In both mineral suites, as Mg-number
decreases, the Cr2O3 content decreases and both TiO2
and Fe2O3 increase even as the Cr-number increases
and may decrease within individual samples (Fig. 14d).
These internal trends demonstrate that some samples
have two or more populations of spinel within each, an
indication of magma mixing.
The populations of olivine and spinel within individual samples are not directly related to the bulk compositions of the rocks. Two samples apiece containing
Group 1 and Group 2 phenocrysts, identified in
Fig. 14c by being enclosed in boxes in the lists of sample
names beneath the data clusters, are among the group
circled in Fig. 3e with the arrow labeled `mixing'
passing through it. Olivine and spinel characteristic
of different stages of differentiation thus mixed
with even more strongly differentiated magma. The
strongest contrast between a primitive (Group 1)
phenocryst assemblage and both the composition and
proportion of the differentiated magma into which the
phenocrysts were mixed is sample PF-05 (* in Fig. 14c;
identified in Fig. 3e).
The compositions of olivine and spinel provide
important information on the conditions in which volatiles were trapped within them. The compositions of
either mineral, for example, can be used to estimate the
MgL [ ˆ Mg/(Mg ‡ Fe2‡)] of the liquids from which
they crystallized (Roeder & Emslie, 1970; Allan et al.,
1988; Allan, 1992). Figure 14f compares the results of
such calculations for samples from Juan Fernandez
and Kilauea±Puna Ridge using all samples for which
coexisting olivine and spinel compositions have been
determined. At both places, the spinel that crystallized
from the most iron-rich liquids is titanian chromite,
with 47% TiO2.
Ideally, the two calculations should provide nearly
the same estimated MgL, and indeed a regression
through the data for Kilauea and Puna Ridges falls
very nearly on a 1:1 trend. This is not the case for Juan
Fernandez: there is a divergence with decreasing MgL,
such that spinel appears to have crystallized from more
iron-rich melts than immediately intergrown olivine as
differentiation proceeded. In detail, a number of
samples from Kilauea and Puna Ridge show the same
divergence, but these are not sufficient to shift the
regression very much.
To understand this, some of the assumptions built
into the calculations need to be considered. In the first
place, to determine MgL, the value of the Fe±Mg
partition coefficient for olivine, KD, must be assumed,
even for spinel, as the calculated exchange free energy
for spinel, G EX, depends on it (Sack, 1982; Sack &
Ghiorsio, 1991). It is often taken to be 030, as originally determined by Roeder & Emslie (1970), but it can
also be calculated from glass compositions using an
algorithm of Carmichael & Ghiorsio (1990). For
Kilauea±Puna Ridge glasses with MgL563, this
value indeed is appropriate. I use the same assumption
for all Juan Fernandez olivine except the three most
magnesian in Fig. 14f. Above MgL ˆ 63, the algorithm of Sack & Ghiorsio (1991) indicates that KD
should increase to 034 among the most magnesian
glasses from Kilauea and Puna Ridge (see Wilkinson &
Hensel, 1988), and I have taken this into account for
those glasses. For the three samples with most magnesian olivine from Juan Fernandez, for which no glass
compositions exist, I interpolated using a linear regression between olivine compositions and Puna Ridge
glasses more magnesian than MgL ˆ 63. Regardless
of complications for primitive compositions, however,
Juan Fernandez samples diverge from those of Kilauea±
Puna Ridge most strongly below MgL 5 60, and among
these, KD ˆ 030 is appropriate for both places.
A second assumption concerns the MgL of the basaltic liquid from which the minerals crystallized. It is
related to the oxidation state of that liquid, and will
vary depending on the ratio Fe2‡/(Fe2‡ ‡ Fe3‡) of the
melt. For glasses analyzed by electron microprobe, this
ratio has to be assumed. The data plotted in Fig. 14f
are based on Fe2‡/(Fe2‡ ‡ Fe3‡) ˆ 086, which is a
common enough assumption, for example, in calculating CIPW norms. This ratio, however, is somewhat
higher than has been determined for many abyssal
tholeiites, for which an average value of at least 090
is recommended (Christie et al., 1987). It is also not
likely to be appropriate across the board for either
Kilauea±Puna Ridge (Carmichael & Ghiorsio, 1990)
or Juan Fernandez tholeiites. The assumption of a
constant Fe2‡/(Fe2‡ ‡ Fe3‡), however, when applied
to estimation of MgL using coexisting olivine and spinel, can indicate differences in oxidation state of host
liquids. Higher Fe2‡/(Fe2‡ ‡ Fe3‡) than 086, for
example, will shift MgL calculated from olivine to
lower values, more in accord with those calculated
from spinel. Indeed, Juan Fernandez spinel compositions have systematically lower Fe-number [ ˆ Fe3‡/
[(Fe3‡ ‡ Cr ‡ Al)] at any given Mg-number than
spinel from Kilauea±Puna Ridge. On this basis, the
Juan Fernandez tholeiites were not as oxidized during
crystallization of their phenocrysts, especially those
with more differentiated compositions.
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Departures from an ideal 1:1 correspondence
between MgL values calculated for olivine and spinel
might be a consequence of partial re-equilibration of
spinel with melt before entrainment in current host
lavas. Scowen et al. (1991) showed that an initial population of spinel at Kilauea Iki's lava lake became richer
in Fe2‡, Fe3‡, and Ti, and poorer in Mg, Al, and Cr,
than the liquidus chromite as the lake cooled over
22 years. The transformation resulted both from cooling and a shift in f(O2) as a result of interaction with
the atmosphere and rainwater percolating down
through cracks in the crust in the lava lake. With a
few exceptions, Puna Ridge spinel, which Clague et al.
(1995) viewed to have become entrained together with
associated olivine in more differentiated, iron-rich
magmas, did not experience similar changes in composition. However, although compositions of coexisting
olivine and spinel provided by Scowen et al. (1991)
show a shift toward re-equilibration with more
iron-rich liquids through time (Fig. 14f ), there was
not much of a shift in the relative oxidation state
(with respect to, say, the nickel±nickel oxide buffer)
that could have induced a marked trend away from the
Kilauea±Puna Ridge regression and the ideal 1:1 estimate for melt MgL values. This explanation therefore
appears to be inadequate to explain such a shift
among the Juan Fernandez picrites, leaving change
in oxidation state as the best explanation for these
relationships.
To explain a 5 mol % difference in MgL calculated
from olivine and spinel for a differentiated Kilauea
tholeiite, about half the iron usually present as Fe2O3
should be converted to FeO, producing a 35±4 mol %
increase in the Fo content of olivine, and a reduction of
about 15 log units in f(O2). The matter is complicated
because usually the olivine and spinel did not crystallize from a melt having the composition of the host
basalt, and because the oxidation state may have continued to vary after the minerals crystallized. A
mechanism for reduction in f(O2) was outlined by
Anderson & Wright (1972), who showed that the composition of coexisting ilmenite and magnetite in the
groundmasses of some Kilauea tholeiites is a consequence of reduction in f(O2) after `effervescence' of
the magma during transport and flow as lava. Of
particular importance is the loss of sulfur by the reaction FeS ‡ 3Fe2O3 ˆ 7FeO ‡ SO2. Loss of 500 ppm of
S to degassing of SO2 results in reduction of 072%
Fe2O3 to FeO. None of the Juan Fernandez samples
with Group 2 assemblages of olivine and spinel contains sulfide within inclusions either in olivine or in the
groundmass, and it is present in inclusions only in two
samples with Group 1 phenocrysts. Degassing is therefore implicated as a potential cause for the systematic
divergence during magmatic differentiation of MgL of
Juan Fernandez tholeiitic liquids as estimated using
olivine and spinel. If this is the case, then the vesiculation took place before or at least during crystallization
of olivine and spinel phenocrysts, and therefore before
final mixing of porphyritic magmas with the more
extreme differentiates that resulted in the present
bulk compositions of the rocks. Loss of S by degassing
at Kilauea occurs at pressures mainly below 100 bars
(Moore, 1965; Killingley & Muenow, 1975; Kyser &
O'Neil, 1984), thus much of the history of crystallization and mixing involving olivine and spinel among
Juan Fernandez tholeiites, and consequently of incorporation of volatiles including helium into phenocrysts,
took place in the shallow probably subaerial crust.
There, magmas may well have interacted with altered
rock or groundwater before crystallization of olivine.
This might explain the high proportion of atmospheric
argon in volatiles extracted from olivine in Juan
Fernandez samples (Farley & Craig, 1994).
Olivine compositions can also be used to estimate
crystallization temperatures and, of more interest, differences in temperature of magmas involved in mixing.
For Juan Fernandez tholeiites, I use a modification of
the geothermometer of Helz & Thornber (1987) for
Kilauea Iki's lava lake, which is based on temperature
measurements as the lake cooled that varied linearly
with the MgO content of glasses quenched in cores.
Clague et al. (1995) applied this to Puna Ridge glasses.
Their temperatures can consequently be related to
glass MgL, and this in turn to the values of MgL calculated for olivine and spinel. For Puna Ridge glasses, the
relationship between MgL and temperature breaks
down into two linear trends with different slopes
Fig. 14 (opposite). Summary of compositions of olivine and spinel in Juan Fernandez olivine tholeiites. Sample groups are indicated by
number, and different shading or symbol color. (a) Histogram of olivine compositions with a principal peak at Fo82. (b) Histogram of spinel
Mg-numbers [ ˆ Mg/[Mg ‡ Fe2‡)]. The distribution is bimodal, defining sample groups 1 and 2, comprising the ranges of spinel enclosed in
olivine (black and dark gray, respectively), and (light gray) spinel with lower Mg-number in the groundmasses of the rocks. (c) Spinel Mgnumber vs Fo contents of immediately adjacent intergrown olivine. This figure provides the basis for the groupings in (a) and (b). (d) Spinel
Mg-number vs Cr-number. Filled circles, Juan Fernandez tholeiites, shaded as in (a) and (b); open diamonds and triangles, Kilauea±Puna
Ridge tholeiites and xenocrysts, respectively (D. Clague, unpublished data). The shaded field for spinel in Pacific MORB and seamounts is
based on data of Natland et al. (1983), Natland (1989) and Natland (unpublished), plus Allan et al. (1988, 1989). (e) Mg-number vs Ti for
spinel. Symbols and eastern Pacific array are as in (d), except that a light gray shaded field represents data for Kilauea and Puna Ridge. (f)
MgL calculated from olivine based on the procedure of Roeder & Emslie (1970) vs MgL calculated from coexisting spinel, based on the
procedure of Allan et al. (1988). (See text for details.) Symbols are as in (d), with addition of bold open squares representing re-equilibrated
spinel from Kilauea Iki lava lake (Scowen et al. (1991). Trends JF and KP are respective linear regressions through the data for Juan
Fernandez and Kilauea±Puna Ridge.
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Fig. 15. The Kilauea geothermometer of Helz & Thornber (1987)
modified to use with glass MgL by calibrating to temperatures calculated for glasses from Puna Ridge by Clague et al. (1995). Two linear
regressions shown in the figure can be used, corresponding to olivinecontrolled liquids with MgL 4 60, and MgL 60, where MgL is
calculated using Fe2‡/(Fe2‡ ‡ Fe3‡) ˆ 086.
(Fig. 15), the less steep one being for glasses with MgL
562, corresponding to the transition between olivinecontrolled and multiply saturated glasses. Other
geothermometers, such as that of Montierth et al.
(1995) for Mauna Loa, give different temperatures by
about 20±40 but the same differences in temperatures
of contrasting magmas involved in mixing.
Ranges and averages of crystallization temperatures
of olivine phenocrysts, olivine rims and groundmass,
and spinel, in Juan Fernandez tholeiites are presented
in Fig. 16. Samples are listed downward in order of
decreasing average olivine phenocryst temperature (or
Fo content). The bar plot to the right shows the difference in temperatures calculated for average phenocryst
and average rim or groundmass olivine, and the graph
at the bottom gives both Fo content and MgL as functions of temperature, with the breaks in slope corresponding to the transition between olivine-controlled
and cotectic crystallization of olivine, clinopyroxene,
and plagioclase. This took place at 1150 C.
The ranges of temperatures calculated for the spinel
populations shown in Fig. 16 are similar to the ranges
found for rim or groundmass olivine, not large olivine
phenocrysts. However, as discussed above, this apparently is the result of too high an estimate for the temperatures of olivine crystallization on the assumption of
constant Fe2‡/(Fe2‡ ‡ Fe3‡). If this ratio increased by
degassing of SO2, then temperatures of olivine crystallization were lower, and those of spinel and olivine
phenocrysts would correspond. The actual temperatures of crystallization of olivine phenocrysts in sample
Fig. 16. Calculated average temperatures and temperature ranges of
crystallization for olivine and spinel populations in Juan Fernandez
olivine tholeiites and alkalic olivine basalts, arranged in order of
decreasing temperatures. Dashed lines give ranges to individual
extreme estimated temperatures. The bar graph to the right gives
differences between average temperatures calculated for olivine
phenocrysts and olivine rims plus groundmass crystals. Actual olivine
compositions and calculated equilibrium MgL values can be interpolated from the curves at the bottom.
PIN-12, for example, were probably less by 50 , and
thus closer to 1200±1150 C. Groundmass olivine
crystallized below 1150 C.
Temperature contrasts between average phenocrystic and rim or groundmass olivine range between 20
and 60 C. These represent the differences in average
temperature of crystallization of the two populations of
olivine in each sample, and record the fundamental
information about mixing of magmas. The cooler magmas involved in mixing in each case were at or below
1150 C. In four cases based on spinel the hotter magmas were just above this temperature, and in three
cases they were at or below it. These estimates support
the earlier inference from bulk compositions that mixing took place between primitive, olivine-charged
magma and cooler, significantly more differentiated
multiply saturated magma. However, apart from very
rare individual skeletal plagioclase phenocrysts, there
is no indication in any of the samples that silicate
minerals other than olivine were crystallizing at the
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HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
time that mixing occurred, nor, indeed, until the
groundmasses crystallized after eruption. This may
have been because the differentiated mixing endmember was usually nearly aphyric, or because such
phenocrysts as it contained were resorbed upon mixing
with hotter magma. After mixing, the hybrid magmas
were probably no longer on a multiphase cotectic,
and both olivine and spinel continued to crystallize
by themselves at temperatures below 1150 C.
DISCUSSION
Volatiles including helium were captured by olivine
phenocrysts in tholeiitic basalts and picrites from the
Juan Fernandez Islands. The volatiles occur in cavities
produced when olivine managed to grow around portions of bubbles attached mainly to discontinuities on
their surfaces, and in other cavities arranged along
secondary fractures in those olivine crystals. In both
cases, the cavities exist because the host magma was
losing volatiles by formation and escape of bubbles at
the time of crystallization. In several cases, significant
degassing of SO2 had already occurred at low pressure
before most olivine phenocrysts formed, reducing the
oxidation state of host liquids, with this influencing the
compositions of both olivine and spinel. Almost all
olivine is phenocrystic; virtually none was entrained
from dunite cumulates or frozen porphyritic basalts.
Mixing occurred at two stages, the first as different
populations of olivine and spinel aggregated, and the
second when those aggregates mixed with strongly
differentiated basalt magma. Vesiculation and volatile
capture accelerated when magmas were stirred or
mixed.
Dixon et al. (1991) and Clague et al. (1995) presented
a model for Kilauea and Puna Ridge, the basalts of
which resemble those of Juan Fernandez in so many
ways, involving mixing between degassed differentiated magma developed at high levels in the crust,
and usually vapor-saturated porphyritic magmas at
depth. The model presumes that the degassed therefore
dense shallow magmas sink through a long (4 km)
magma column, and that this enabled mixing with
olivine tholeiite. Although the model for volatiles does
not depend on the origin of olivine phenocrysts in the
picrites, Clague & Denlinger (1994) and Clague et al.
(1995) preferred an origin for many of these crystals,
namely those with blocky morphologies and containing subgrains, by scavenging from partially deformed
dunite cumulates, present in the deep Kilauea crust.
For Juan Fernandez tholeiites, on the other hand,
most olivine crystallized at depths shallow enough
to have already experienced significant degassing,
particularly of SO2, which affected oxidation state.
I propose that the evidence for deformation in some
crystals, most of which are relics of euhedral minerals
rather than dunite cumulates, resulted from nonhydrostatic forces experienced during magma transport. Komar (1972, 1976) showed that when olivine
phenocrysts reach concentrations of only 8%, they
begin to interfere with each other during magma transport, with the consequent flow differentiation resulting
in very high concentrations of phenocrysts within, for
example, the centers of dikes (see Drever & Johnston,
1958). At high phenocryst concentrations, during flow
through irregular fractures locally having narrow orifices, olivine is bound to collide with other olivine as
well as the walls of the conduit. Pulsing of magma will
produce variable flow rates. During lulls in the flow,
the olivine will settle onto ledges and discontinuities in
the conduit. Renewed high flux and fracturing along
conduit pathways will pulse magma through local
accumulations of crystals, entraining them as it passes.
There is therefore ample opportunity for nonhydrostatic stresses to develop in a magma conduit
system, and thus for deformation features to form in
euhedral olivine either suspended in, or temporarily
settled from, the flowing magma. Both the irregular
outlines of blocky olivine and rounding can result from
collisions of crystals, with the latter being analogous to
tumbling in a lapidary. Mechanical breakup of minerals with subgrains or of olivine with contraction fractures produced those few grains that are completely
angular in outline.
Through all of this, the magmas vesiculated. Vesiculation began at fairly high pressure when magmas
reached saturation with CO2 during their ascent (e.g.
Bottinga & Javoy, 1990; Dixon et al., 1991). However,
most vesiculation at Juan Fernandez occurred as a
consequence of two processes: stirring in the presence
of olivine phenocrysts, and magma mixing. This is
because supersaturation is a metastable state during
which bubbles form only after the activation energy
for their nucleation is overcome. Carbonated water in
a glass may effervesce slowly, or not at all, but it might
foam over the rim if it is stirred, or if ice is dropped into
it. Later, even when it is comparatively flat, stirring
will cause bubbles to nucleate on the floating ice. Thus
the movement of magma will accelerate vesiculation,
as will a sudden influx of picrite into a nearly stagnant
pool of partially degassed ferrobasalt magma. Besides
providing olivine on which bubbles can nucleate, the
influx will also add a certain amount of superheat
to the ferrobasalt. This by itself will accelerate vesiculation. A similar effect evidently triggered certain
eruptions on Iceland. There, addition of superheat
by influx of basalt into the base of a zoned magma
chamber at Askja volcano produced convection,
intense vesiculation, and explosive eruption of rhyolite
floating at the top of the chamber (Sparks et al., 1977).
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Figure 16 illustrates that mixing at Juan Fernandez
was between magmas different in temperature by
20±60 C. This is far less than the temperature contrast
between basalt and rhyolite, thus an explosive consequence to mixing probably would not result. However,
convective stirring and an increase in the rate of
vesiculation seem likely.
What was the origin of secondary inclusions along
annealed curvilinear fractures and the material in
them? I presume again that the fractures resulted
from non-hydrostatic forces acting on phenocrysts,
but the absence of consistent orientations for the trains
of inclusions, indeed the bewildering diversity of orientations of these trains in individual crystals, does not
seem consistent with the effects of a uniaxial compressive stress or of shear. Instead, I propose that these are
contraction fractures, induced by the sudden introduction of host magma into cooler differentiated basalt
during mixing. The porphyritic host was already vesiculating because of stirring, and bubbles were already
attached to the growing olivine. An analogy is the
dropping of ice into carbonated water. The temperature contrast is only a few degrees, but cracking of the
ice is audible, and one is usually hard pressed to keep
the liquid from foaming over the top of the glass. A
network of cracks generally forms in each cube of ice.
Ice, of course, is one of the few solids that expand
upon cooling across the phase transition from liquid.
This happens when warm water is poured over it, and
the cracks that form are a consequence of the nonuniformity of this expansion. Olivine, on the other
hand, will contract when suddenly immersed in cooler
melt, thus opening up fracture porosity into which
surrounding boundary-layer melt can be drawn. The
boundary layer at this stage consists of a heterogeneous
assemblage of silicate melt with strong chemical gradients in certain components, tiny spinel, bubbles, and,
depending on the pressure, immiscible sulfide droplets.
The three fluid phases have different viscosities and
compressibilities, thus they will be prone to separate
physically as they pass quickly into narrow fractures. A
large ratio of mineral surface area to melt in the fractures promotes quick annealing of the fractures, with
one consequence being enrichment in iron of the
residual silicate melt in the fractures. Another is to
decrease the proportion of trapped silicate melt with
respect to that of spinel, bubbles, and sulfides in the
fractures.
All the while, primary inclusions are captured in
their usual fashion as the growing crystals respond to
heightened undercooling, development of boundary
layers, and local constitutional supercooling near bubbles on crystal faces. Primary and secondary inclusions
form at almost exactly the same time, as the crystals are
suspended in vesiculating, sometimes frothing, melts.
NUMBER 3
MARCH 2003
In both cases, many bubbles are already in the melt.
They both nucleate on and stream around the growing
olivine. In both cases, much volatile capture is closely
tied to magma stirring and mixing. The mixing
heightens undercooling, accelerates vesiculation, and
promotes thermal cracking of the olivine, which draws
melt laden with dissolved and exsolved volatiles into
the crystal interiors.
Very little olivine was incorporated from dunite in
this process. In dunite cumulates, olivine grains are
anhedral and irregular in shape, as already mentioned,
whereas almost all olivine in Juan Fernandez tholeiites
is faceted and euhedral. Only a single dunite xenolith
was found in one sample (MF-20), and it was derived
from a substantially differentiated liquid (Fo826,
MgL ˆ 57). Some olivineÐperhaps certain of the
grains with broken edgesÐmay have been derived from
loose crystal aggregates (Schwindinger & Anderson,
1989), but not dunite. On the other hand, settling of
olivine and spinel from these tholeiites would produce
dunitic cumulates with a wide range of olivine compositions. The cumulates themselves would be laden with
volatiles, many of the grains will have curvilinear trains
of secondary inclusions in miscellaneous orientations,
and probably about 10% of them will have subgrains
and planar deformation lamellae. Therefore there is no
reason to suppose that any of these features in dunite
are necessarily the result of deformation or of reaction
with percolating magma (e.g. Kirby & Green, 1980).
Clague & Denlinger (1994) proposed that there is a
substantial body of dunite deep in the volcanic structure of Kilauea volcano. They argued that this is a
currently deforming mass, that the deformation is
responsible for episodic lateral failure of the flank of
the volcano, and that olivine-controlled basalts from
Kilauea and Puna Ridge derive much of their olivine,
particularly that with kink bands, from the dunite.
There is no evidence bearing on the existence of similar
masses at Juan Fernandez. However, the high concentrations of olivine phenocrysts in some Juan Fernandez
tholeiitic picrites probably resulted from flowage differentiation in narrow dikes that reached to very high
levels in the volcanic structure. There are ample means
for olivine to deform during magma transport in narrow dikes. Much of the olivine crystallized at depths
shallow enough so that significant degassing of SO2
and resultant reduction of the magmas had already
taken place. In some cases, it crystallized from liquids
100 or more cooler than postulated primitive Kilauea
liquids. The magmas containing the olivine mixed with
differentiates that were 20±60 cooler still. Therefore
upward and lateral movement of the picritic material
along rift zones must force it into contact with
high-level differentiates where mixing can took place,
and thus olivine-rich cumulates should be distributed
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NATLAND
HELIUM CAPTURE BY OLIVINE PHENOCRYSTS
at many levels in Juan Fernandez volcanoes. There is
no evidence that degassed differentiates sank for several
kilometers to reach the picrites. Finally, mantlederived helium persisted as a constituent among the
volatiles of these magmas to points far along rift zones
and away from any central conduits. By analogy to
Kilauea, these probably functioned at depths only
some 2±4 km beneath the surface at any one time,
and rose with the volcanoes as they grew (Ryan, 1987).
The processes described here must be very common.
Phenocrysts in most ocean island basalts studied so far
contain volatiles in inclusions. They are trapped in
bubbles and fractures and are extractable by simple
crushing. Craig et al. (1993) found rare gases in phenocrysts from 21 ocean island basalts, with helium isotope
ratios ranging from MORB or sub-MORB values to
`High-3He' ratios similar to the Juan Fernandez
basalts. They also found evidence for a common atmospheric rare-gas component in these phenocrysts,
designated `ASW' (atmospheric or seawater gases),
which were present in magma chambers during growth
of olivine and clinopyroxene, followed by extensive
degassing. Because the ASW component resides within
the phenocrysts, it must have been derived directly from
fresh water, from seawater, or from fresh- or seawatersaturated lithosphere adjacent to magma chambers.
Craig et al. (1993) proposed that introduction of extraneous CO2 from the ASW source triggered supersaturation in the magma and the production of bubbles
that were then trapped in the growing phenocrysts.
The acceleration of vesiculation by means of mixing,
as proposed here, however, is fundamentally indifferent to how volatiles entered the magmas or the
attributes of their sources. If a magma is already supersaturated in volatiles to begin with (Bottinga & Javoy,
1990), perhaps in a holding reservoir at a shallow level
in the conduit, but just below sea level, then a simple
perturbation in pressure or temperature causing degassing and change in oxidation state could induce some of
the same effects attributed here to mixing, including
zoning of minerals (Anderson & Wright, 1972). The
ASW source would be evident either way. However,
the usual way that magmas move to higher levels in a
conduit or to eruption is by displacement by other
magmas introduced from deeper in the Earth. Mixing
in conduits or rift zones and vesiculation are both
extremely common processes. The two acting in combination are a major mechanism for incorporation of
noble gases into phenocrysts.
CONCLUSIONS
Mantle-derived helium, other noble gases, and more
common volatile species such as CO2, S, and H2O,
were incorporated into olivine phenocrysts in basalts
and picrites from the Juan Fernandez Islands primarily during episodes of mixing between differentiated
and primitive lavas at depths sufficiently shallow in
the volcanoes' structure for magmas to have lost significant SO2, reducing the melt oxidation state in the
process. The mixing and its accompanying turbulence
accelerated vesiculation, with much bubble nucleation
occurring on numerous seed crystals of olivine. A great
variety of primary multiphase inclusions, containing
silicate melt, Cr-spinel, immiscible sulfide melt, and
bubbles, the last three of which tended to nucleate on
the olivine crystal surfaces, were captured at the heightened undercooling induced by both vesiculation and
mixing. Some olivine also experienced thermal shock
during the mixing, resulting in the formation of secondary curvilinear contraction fractures into which
boundary-layer melts were nearly instantaneously
drawn, thence to become trapped along the fractures.
Compositions of spinel and olivine indicate that
mixing occurred between less and more differentiated
magmas, some 20±60 different in temperature, and
accordingly in bulk composition, but also in oxidation
state, with the more differentiated mixing components
being less oxidized. Crystallizing olivine was consequently more iron rich than it would otherwise have
been. Incorporation of atmospheric argon may have
occurred at this time.
Considering all the processes involved, departure
from equilibrium was the norm. Degassing, mixing,
reduction in oxidation state, supercooling resulting
from vesiculation, constitutional supercooling, preferential nucleation of various phases at locations determined by interface instabilities on olivine crystal
surfaces, physical separation of solid, liquid, and
vapor phases in contraction fracturesÐthese are all
rate-controlled rather than equilibrium processes.
They complicate the potential to reconstruct equilibrium phase relationships and compositions, as well
as extensive and intensive conditions of crystallization,
using any suite of inclusions.
Mantle-derived volatiles including helium occur in
olivine from every sample studied. Such gases clearly
pervaded magma conduits and storage reservoirs,
reaching shallow levels in the structure of Juan
Fernandez volcanoes, including rift zones where magmatic differentiation, mixing, and degassing were all
extensive.
ACKNOWLEDGMENTS
I especially thank Harmon Craig, who suggested this
study, encouraged it at all stages, helped interpret
many of the micrographs, explained to me the intricacies of the isotope geochemistry of noble gases,
and argued many points pro and con throughout the
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JOURNAL OF PETROLOGY
VOLUME 44
preparation of the manuscript. Harmon and Ken Farley collected the Juan Fernandez picrites on Scripps
Institution of Oceanography Expedition Hydros with
the assistance in the field of Ron Comer and Valerie
Craig, and the co-operation and help of the officers and
crew of R.V. Melville. I thank Ken Farley for the thin
sections, the grain mounts, the olivine separates, the
prior geochemical study, and continuing helpful discussion. Roy Fujita provided able assistance with the
electron microprobe, as did Patricia Blackwelder and
Teri Hood with the scanning electron microscope. The
University of Miami supported establishment of the
Petrography Laboratory at the Rosenstiel School,
which was largely designed based on the requirements
of this work. Ken Farley, Dave Clague and Richard
Arculus reviewed the manuscript, to its considerable
benefit. The Mantle Geochemistry and Volcanology
Program of the Earth Sciences Division, US National
Science Foundation, supported this work.
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