Akizawa et al.
Progress in Earth and Planetary Science(2024) 11:38
https://doi.org/10.1186/s40645-024-00643-w

PAPER WITH FULL DATA ATTACHED

Progress in Earth andPlanetary Science

Open Access

Evidence for suboceanic small‑scaleconvection from a “garnet”‑bearing lherzolite
xenolith from Aitutaki Island, Cook Islands
Norikatsu Akizawa1,2* , Kazuhito Ozawa1, Tetsu Kogiso3, Akira Ishikawa4, Akira Miyake5, Yohei Igami5,
Simon R. Wallis6, Takayoshi Nagaya2,6, Chihiro Ohshima1, Ryo Fujita4, Tatsuhiko Kawamoto7, Akihiro Tamura8,
Tomoaki Morishita8, Shoji Arai8 and Atsushi Yasumoto5

Abstract
Garnet peridotite xenoliths have been rarely reported from suboceanic mantle. Petrographic and geochemical
characteristics of garnet-bearing oceanic peridotite xenoliths provide precious information on dynamics of the suboceanic lithosphere and asthenosphere interaction. We examined a lherzolite xenolith included in olivine nephelinite
lava from Aitutaki Island, a member of the Cook-Austral volcanic chain. The lherzolite xenolith contains reddish finegrained (< 5 µm in size) mineral aggregates (FMAs) with size range of 0.5–6 mm, consisting of olivine, calcic and sodic
plagioclases, aluminous spinel, native iron, and nepheline. Microstructural observations and chemical data corroborate that the FMA is a decomposed pyrope-rich garnet including chromian spinel grains with an irregular highly
indented morphology in the center. The FMA is surrounded by pyroxene-poor and olivine-rich aureole. The spatial
and morphological relationships of FMA and chromian spinel with pyroxene-depleted margin suggest a reactionof aluminous spinel + pyroxenes → pyrope-rich garnet + olivine, which requires a compression before decompositionof the garnet to FMA. An orthopyroxene grain shows slight but clear chemical zoning characterized by increase in Al,
Ca, and Cr from the grain center to the rim. The zoning patterns of Al and Ca in the orthopyroxene grain can be modeled by diffusion-controlled solid-state reactions induced by pressure and temperature changes, keeping surface concentrations in equilibrium with the other coexisting mineral phases. The results indicate that the mantle, from which
the lherzolite xenolith was derived, underwent isothermal decompression followed by a weak heating on a time scaleof a few tenths of million years before the xenolith extraction. From the deduced compression and decompression
histories, we hypothesize that the mantle beneath Aitutaki Island was once dragged down to a garnet-stable deep
mantle region and brought up later by small-scale sublithospheric convection.
Keywords Garnet lherzolite, Oceanic lithosphere, Native iron, Mineral diffusion, Liquid immiscibility, Carbonaceous
fluid, Nephelinite*Correspondence:
Norikatsu Akizawa
akizawa@g.ecc.u-tokyo.ac.jp
Full list of author information is available at the end of the article© The Author(s) 2024. Open Access This article is licensed under a Creative Commons Attribution 4.0 International License, which
permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give appropriate credit to theoriginal author(s) and the source, provide a link to the Creative Commons licence, and indicate if changes were made. The images orother third party material in this article are included in the article’s Creative Commons licence, unless indicated otherwise in a credit line
to the material. If material is not included in the article’s Creative Commons licence and your intended use is not permitted by statutory
regulation or exceeds the permitted use, you will need to obtain permission directly from the copyright holder. To view a copy of thislicence, visit http://creativecommons.org/licenses/by/4.0/.

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

1 Introduction
Garnet peridotite xenoliths are of significant importance
to extract information of deep lithospheric and asthenospheric mantle as they are derived from the depth deeper
than ~ 70 km (O’Hara et al. 1971; Klemme and O’Neill
2000). The garnets in the peridotites provide useful andreliable geothermobarometry indispensable for reconstructing chemical and thermal structure of the deep
lithospheric mantle (Nickel and Green 1985; Smith andBarron 1991; Griffin et al. 1999a, b; Ishikawa et al. 2004).
The garnet peridotite xenoliths have been frequently
reported in kimberlites mostly occurring in the cratonic
region (Carswell and Dawson 1970; Boyd and Nixon
1978; Richardson et al. 1985), whereas their reports fromthe oceanic region are rare.
Many garnet-bearing xenoliths of pyroxenite and wehrlite have been reported from Hawaii Islands (e.g., Sen
et al. 2005), but a garnet peridotite xenolith is reported
only in Oahu Island by Goto and Yokoyama (1988) without unequivocal evidence. Malaita Island in Solomon
Islands is one of the rare localities yielding garnet peridotite xenoliths, which have contributed to reconstructchemical and rheological structure of the deep lithospheric mantle beneath Ontong Java Plateau (Ishikawa
et al. 2004; Ishikawa et al. 2011; Tommasi and Ishikawa
2014; Demouchy et al. 2015). Harigane et al. (2011)
reported a fine-grained aggregate of spinel and orthopyroxene in a peridotite xenolith from petit-spot volcano
in the northwestern Pacific Ocean. They argued that thefine-grained aggregate is of garnet origin without providing detailed petrography and mineral chemistry.
These limited studies on the oceanic garnet peridotite are insufficient to illustrate the spatial and temporalvariation of the oceanic lithosphere, which is imperative
to understand the evolution of lithosphere and asthenosphere boundary beneath the oceans (e.g., Fischer et al.
2010; Kawakatsu and Utada 2017; Rychert et al. 2020).
Thus, we need to expand the localities and samples ofoceanic garnet peridotite, and to examine their detailed
petrological and geochemical characteristics. To achieve
this purpose, it is crucial to look for the garnets in theperidotites, but they are in general susceptible to low
pressure breakdown reaction during xenolith transportation from the mantle to the surface. Hence, reaction
product after the garnet is a fundamental study target
when supported by careful and exhaustive investigations
to verify decomposition process.
Here, we present a peridotite xenolith from Aitutaki
Island, a member of the Cook-Austral volcanic chain,
where many peridotite xenoliths have been studied(Fodor et al. 1982; Snortum et al. 2019). The peridotite
xenoliths contain aggregates of very fine-grained minerals, which were interpreted to be a breakdown product ofPage 2 of 28

aluminous orthopyroxene in the previous studies (Fodor
et al. 1982; Snortum et al. 2019). We show that they area breakdown product of garnet in fertile peridotite withseveral lines of evidence from microstructural, mineralogical, and geochemical data. We then characterize
the fertile garnet peridotite by determining whole-rockelemental concentrations, including highly siderophileelements (HSE) and Re–Os isotope ratio. We further
clarify an ascent history of the peridotite xenolith by
reproducing chemical zonings of orthopyroxene andmicrostructures related to reactions induced by pressure–temperature changes and fluid influxes. Combining
all these results, we eventually infer the mantle dynamics
beneath the South Pacific Ocean lithosphere.

2 Geological background
Aitutaki Island is an 8 km × 3 km volcanic island with anatoll lagoon located at the northwest end of the CookAustral volcanic chain in the Southwest Pacific Ocean(Fig. 1a). The island chain is oriented in northwest-southeast extending over ca. 2000 km. The Cook-Austral volcanic chain is composed of 13 islands: Marotiri, Rapa,
Raivavae, Tubuai, Rurutu, Rimatara, Mangaia, Mauke,
Mitiaro, Atiu, Rarotonga, and Aitutaki Islands (Fig. 1a).
This area has been interpreted to be a superswell hosting large thermochemical anomaly in terms of seismic
tomography and anomalously shallow seafloor depths
of a few hundreds of meters (McNutt and Judge 1990;
Adam and Bonneville 2005). This superswell is presumed
to have formed above a superplume from the deep mantle. The Cook-Austral volcanic chain is located above theskirt of the large low shear velocity province (LLSVP) in
the deep mantle (Torsvik et al. 2006).
The volcanic rocks of the Cook-Austral Islands weredated using K–Ar and Ar–Ar geochronometries (e.g.,
Duncan and McDougall 1976; Turner and Jarrard 1982;
Bonneville et al. 2006; Hanyu et al. 2013; Rose and Koppers 2019). Recent determination of the volcanic ages
with eliminating those affected by weathering and alteration revealed two distinctive trends in distance-age plot.
They are called Macdonald trend (Marotiri + Rapa + Raivavae + Tubuai + Rurutu + Rimatara + Mangaia Islands,
ages ranging from 4 to 20 Ma) and Rurutu trend(Rurutu + Mauke + Atiu + Aitutaki Islands, ages ranging
from 1 to 9 Ma) (Rose and Koppers 2019). At the locations extending the two trends to 0 age, there are currently active Macdonald seamount and Arago seamountthat was active at ~ 0.23 Ma (Bonneville et al. 2002)(Fig. 1a).
Aitutaki Island is situated on the ~ 107 Ma oceanic
lithosphere according to the global plate model by Müller et al. (2019) (Fig. 1b). The volcanic activities of
Rurutu and Aitutaki Islands show two age peaks: ~ 13 Ma

Akizawa et al. Progress in Earth and Planetary Science(a)(2024) 11:38

Page 3 of 28

Old Aitutaki (~10 Ma)

3000 mAtiu Mitiaro
Mauke

0m

Young Aitutaki(~2 Ma)Rarotonga
Mangaia-3000 mYoung Rurutu(~1 Ma) Arago SM

20°S-6000 mRimatara Old Rurutu (~13 Ma)
Tubuai

18°S

22°S

Raivavae
24°S

Rapa

26°S

Marotiri
28°S

Macdonald SM
162°W

158°W

154°W

150°W
18°S

(b)

142°W

(c)

20190318A-11-27
AI70

109

107

146°W

Aitutaki

Mitiaro

105

20°S

Atiu

103

Mauke

101

Rarotonga
99−160°W

Mangaia
97

95

0

5 km

22°S

−158°W

Fig. 1 Maps indicating Aitutaki Island and sample localities. a Magnified area of the Pacific Plate. Rectangle inset shows its wide view. The islands
pointed by the orange arrows are classified as Macdonald trend, those pointed by green arrows are classified as Rurutu trend, and those pointed
by the red arrows are classified as Rarotonga trend. Volcanic rocks from Mitiaro are severely weathered and no reliable age data are available (Turnerand Jarrard 1982; Rose and Koppers 2019). b Seafloor age around Aitutaki Island. c Sample locations on Aitutaki Island. The images in (a) werecreated with GeoMapApp application (Ryan et al. 2009). The seafloor age-gridded data presented in (b) are from Müller et al. (2019). The island mapin (c) is after Wood (1978). SM in (a) represents seamountand ~ 1 Ma in Rurutu Island (called old Rurutu stageand young Rurutu stage, respectively), and ~ 9.4 Ma
and ~ 1.4 Ma in Aitutaki Island (called old Aitutaki stageand young Aitutaki stage, respectively). In Aitutaki
Island, many peridotite xenoliths are included in olivine
nephelinite grouped in the young Aitutaki stage (Fodor

et al. 1982; Snortum et al. 2019). The whole-rock chemical composition of olivine nephelinite is marked by hightotal alkali contents ­(Na2O + ­K2O > 5 wt%) (Snortum et al.
2019).
The size of peridotite xenoliths from Aitutaki Island
ranges from a few centimeters to 10 cm in diameter,

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

and they are lherzolite, harzburgite, or dunite (Fodor
et al. 1982; Snortum et al. 2019). We searched xenoliths
at a northern locality (18.83216° S and 159.77441° W),
where olivine nephelinite lavas occur (Fig. 1c). The lavas
include mantle xenoliths with megacrysts of clinopyroxene and ilmenite with ~ 5 cm in diameter. A lherzolite
xenolith sample AI70 examined in detail in this study was
included in an olivine nephelinite sample 20190318A11-27 (Fig. 1c) with other fertile lherzolite xenoliths. The
size of AI70 lherzolite is 3 cm in diameter. The olivine
nephelinite consists of groundmass of clinopyroxene,
nepheline, and magnetite with minor phenocrysts of olivine and clinopyroxene (Wood 1978).

3 Petrography
We examined a thin section of AI70 lherzolite having a
direct contact with the host olivine nephelinite, which
occupies ~ 10% of the entire thin section area (Fig. 2a, b).
The AI70 lherzolite consists mostly of anhedral grainsof olivine, orthopyroxene, clinopyroxene, and finegrained mineral aggregate, which is called FMA hereafter (Fig. 2c). The FMA is reddish purple with the naked
eye and dark red under plane-polarized transmitted light.
A minor amount of spinel occurs in the AI70 lherzolite.
The spinel is dark reddish brown under plane-polarized
light and occurs always included in FMA grains (Fig. 2c;
Supplementary Fig. S1a–d). Detailed petrography of the
FMA and spinel will be described in the next section. The
modal composition of AI70 lherzolite was determinedfrom X-ray intensity maps obtained with electron probe
micro analyzer (EPMA) by using GeoMAC application(https://​github.​com/​Kazuh​ideMi​mura/​GeoMAC). It is
52.0 vol.% olivine, 22.6 vol.% orthopyroxene, 12.8 vol.%clinopyroxene, 11.7 vol.% FMA, and 0.9 vol.% spinel.
Sizes of olivine, orthopyroxene, and clinopyroxene range
respectively 0.1–3 mm, 0.1–4 mm, and 0.1–3 mm, which
are in the ranges of the other Aitutaki lherzolite xenoliths
in the olivine nephelinites. The size of FMA ranges from
0.5 to 6 mm (Fig. 2c; Supplementary Fig. S1d). The AI70
lherzolite is characterized by coarse-grained tabular equigranular texture without undulouse extinction in olivine
and has distinct foliation and lineation defined by shapepreferred orientation of individual grains of pyroxenes
and olivine as well as aggregates of pyroxenes (Fig. 2).
Individual FMA grains have rounded to subrounded outline (Fig. 2). The largest FMA grain is equidimensional(Fig. 2a–c; Supplementary Fig. S1e, f ), and smaller FMA
grains are elongate with long axis parallel to the foliation (Fig. 2c; Supplementary Fig. S1d). These FMA grains
form a cluster, which is elongate parallel to the foliation(Fig. 2c; Supplementary Fig. S1d). The aggregate of FMA
is surrounded by a domain poor in pyroxenes and rich inolivine modes (Fig. 2c; Supplementary Fig. S1d).

Page 4 of 28

Orthopyroxene is locally decomposed into fine-grained(< a few tens of microns in diameter) aggregate, which is
confined to an 1 cm-wide zone from the contact with theolivine nephelinite. Such orthopyroxene grains are recognized as dull brownish color under plane polarized light(Fig. 2a; Supplementary Fig. S2a, b). The decomposed
aggregate consists of clinopyroxene, olivine, richterite,
and glass (Supplementary Fig. S2c, d). Clinopyroxene
occurring in the zone is dull greenish in color underplane polarized light and spongy associated with finegrained (< a few tens of microns in diameter) olivine, spinel, richterite, and glass (Supplementary Fig. S2e, f ). The
grain boundaries in the zone of pyroxene decomposition
are open as wide as 0.3 mm and filled with Na-rich (~ 8
wt% ­Na2O) glass (Fig. 2g; grain boundary sealed withglass, hereafter). Grain boundaries sealed with minerals
are developed outside the zone of pyroxene decomposition, where the grain boundaries are 0.1 to 0.2 mm-wide
and consists of olivine, clinopyroxene, calcite, K-feldspar,
phlogopite, spinel, and rutile with the size up to ~ 100 µm
in diameter (Supplementary Fig. S2h–j; grain boundary
sealed with minerals, hereafter).

4 Detailed petrography and chemistryof spinel‑cored FMA
FMA and enclosed spinel inclusions in AI70 lherzolite
show a peculiar microstructure and heterogeneity inchemical composition. They have records of changes inpressure, temperature, and chemical environment thatthe xenolith sample underwent in the mantle, duringxenolith extraction, and transportation to the surface.
4.1 FMA (Fine‑grained mineral aggregate)

FMA consists of fine-grained polyphase mineral aggregates with the size of individual grains smaller than a few
microns across. Some FMA grains contain spinel grainsin the center (Fig. 2), which will be described in Sect. 4.2.
The FMA has diversity in chemical composition andmicrostructure as seen in Na, K, Ca, Mg, and Fe intensity maps (Fig. 2d–h), from which three types of FMA are
recognized: K-rich, and Na and Ca-poor type (Type 1),
K-poor, and Na and Ca-rich type (Type 2), and heterogeneous type with considerably higher Na than the othertypes (Type 3) (Fig. 2i). Each type has a respective morphology and location of occurrence as follows.
Types 1 and 2 occur in FMA grains more than 1 cm
away from the contact with the host olivine nephelinite (Fig. 2i). Type 1 occurs only in the core region of the
FMA grains, whereas Type 2 occupies a ~ 0.2 mm wide
marginal region with extending toward inside to the coreregion, forming a dendritic network with 0.05–0.1 mmwide branches (Figs. 2i, 3a–b). The main branches havea fracture in the middle sealed with minerals (see below

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 5 of 28

(a)

(e) K
Fig. S
2c
Fig.
S2c

Opx
Ol

Fig. S2e
S2e
Fig.

t
Hos

Fig.
Fig. S2g
S2g

Fig.
3d
Fig. 3
d

FMA

Cpx
Cal
Fig.
Fig. S2h,
S2h, i
10 mm(b)(f) Ca(c)(g) Mg

10 mm

(c)
Ol
Opx

(h) FeCpx
Cr-splFMA
5 mm(d) Na(i) NaType 2Type 1
Type 3Fig. 2 Petrographic characteristics of AI70 lherzolite. a Thin section scan image under plane-polarized transparent light. Almost opaque areas
are FMA, some of which include chromian spinel. The chromian spinel is nearly opaque in background of purple FMA. The positions of Fig. 3d
and Supplementary Fig. S2c, e, g, h, and i are shown with the white squares. b Thin section scan image under cross-polarized transparent light.
The position of c is shown with the white square. c Phase map constructed from the elemental-distribution maps of Fig. 7 using a computational
method (GeoMAC, https://​github.​com/​Kazuh​ideMi​mura/​GeoMAC). d–h Pseudo-color maps of Na, K, Ca, Mg, and Fe distributions, respectively.
Warmer colors indicate higher elemental contents. i Pseudo-color map of Na with type classification of FMA based on petrography and chemicalcomposition (see text for further details). Cal, Calcite; Cpx, Clinopyroxene; Cr-spl, Chromian spinel; FMA, Fine-grained mineral aggregate; Ol, Olivine;
Opx, OrthopyroxeneAkizawa et al. Progress in Earth and Planetary Science(2024) 11:38

for the details). The network of Type 2 decorates isolatedislands of Type 1 with highly irregular morphology in thecore region (Figs. 2i, 3a–b). The Na and Ca contents ofType 1 are lower, and its K content is higher than thoseof Type 2 (Fig. 2d–f ). The abundances of these elements
continuously change between the neighboring two types
over the distance of 0.05–0.1 mm. The minimum Na
and maximum K contents of Type 1 are attained at thecenter of large Type 1 islands (see the pale blue arrows in
Fig. 2i), and such tendencies can be recognized irrespective of the size of Type 1 islands. Small FMA grains arecompletely occupied by Type 2 instead of Type 1 with the
maximum K and minimum Na at the center (see the pinkarrows in Fig. 2i). The Na and Ca contents of Type 2 arehighest in the marginal zone of each FMA grain and tend
to decrease toward the inside (Fig. 2d–f ). These microstructural relationships suggest that Type 2 postdated
Type 1.
Type 3 occurs mainly in FMA grains within ~ 1.4 cm
from the contact with the olivine nephelinite, but it
rarely occurs along the rim of FMA up to ~ 1.6 cm fromthe contact (Fig. 2d–f ). Two smaller FMA grains placed
closer to the contact with the olivine nephelinite (~ 1and ~ 5 mm from the contact, see the green arrows in
Fig. 2i) wholly belong to Type 3, whereas the largest FMA
grain located more than 1 cm away from the contact has
Type 3 partly along the lower left grain margin facing thecontact with the olivine nephelinite (Fig. 2d–f ). The Na
and K contents of Type 3 decrease with distance from thecontact with the olivine nephelinite when the grain margins of FMA are compared.
In each FMA grain, Na is lower, and K is higher at thecore than at the margin (Fig. 2d, e). No systematic variation of Ca content in distance from the contact with theolivine nephelinite is noticed (Fig. 2f ).
All the types of FMA are included in the largest FMA
grain directly contacting with each other (Fig. 3a–i). Type
3 forms a dendritic network with each branch widthPage 6 of 28

thinner than 50 mm and commonly as thin as 20 mm
in the largest FMA (Fig. 3a–d, h). The dendritic morphology emerges from lower left grain margin, which
invades ~ 2.5 mm into the FMA grain and extends for
5 mm along the margin with decreasing the width ofinvasion (Fig. 3a–d). There are differences in morphology of dendritic networks of Types 2 and 3: e.g., thebranch widths are 50–100 µm and 20–50 mm, respectively (Fig. 3a–c). The dendritic networks of Types 2 and3 are mostly complementary with each other, but Type
3 branches locally crosscut those of Type 2 (Fig. 3a–c),
indicating that Type 3 postdated Type 2.
Mineral phases consisting of FMA were identified
by multiple methods (Supplementary Table S1, Supplementary Figs. S3 and S4; see Supplementary methods for analytical details). All the types of FMA consistof olivine, plagioclase, and spinel with or without nativeiron and nepheline. Type 1 consists of polyphase aggregate of olivine (forsterite content: 100 × Mg/(Mg + ­Fe2+)
atomic ratio = 88), calcic plagioclase (anorthite content:
100 × Ca/(Ca + Na + K) atomic ratio = 87), sodic plagioclase (anorthite content = 26), and spinel (Fig. 4) withminor amounts of native iron and nepheline (Fig. 3f ).
Volumetric proportions of the major minerals in Type 1
are estimated to be 36.0 vol% olivine, 9.5 vol% spinel, 25.9
vol% calcic plagioclase, and 28.6 vol% sodic plagioclase(Fig. 4). The size of olivine is 0.5–2 mm, that of calcic plagioclase is 0.2–1 mm, that of sodic plagioclase is ~ 1 mm,
and that of spinel is ~ 0.5 µm across in Type 1 (Figs. 3f
and 4). The size of native iron and nepheline is ~ 5 µm
across and coarser than that of the other main constituent minerals of Type 1. The olivine and calcic plagioclase
are anhedral to subhedral, sodic plagioclase is anhedral,and the spinel is euhedral to subhedral (Fig. 4). The anhedral sodic plagioclase fills interstitial spaces of euhedral
to subhedral grains of olivine, calcic plagioclase, and spinel, forming an interstitial network (Fig. 4). The olivine isin direct contact with both calcic and sodic plagioclases(See figure on next page.)
Fig. 3 Detailed petrographic characteristics of AI70 lherzolite. a–c Pseudo-color maps of largest FMA of Na, K, and Ca distributions, respectively. The
chromian spinel grains have many fractures orienting at high angles against elongate spinel morphology [see the red and yellow arrowheads facing
each other in (c)]. The position of (e) is shown with the white square. d Backscattered electron (COMP) image of chromian spinel grains in largest
FMA. It is noteworthy that native iron grains (tiny white dots) are less abundant around the chromian spinel grains and in the outer margin of FMA.
The positions of (h) and (i) are shown with the white squares. e Magnified view of Fe and Na distribution in largest FMA. Native iron and nepheline
grains are recognized by red spots and pale blue spots, respectively. Note that they tend to be in direct contact with each other. The positionsof (g) and (f) are shown with the white squares. f High magnification COMP image of Type 1 FMA. g High magnification COMP image of Type 2
FMA occupying marginal region around a vein sealed with olivine, plagioclase, and native iron. h High magnification COMP image of Na-rich outerrind of largest FMA close to the contact with the host olivine nephelinite. The position of i is shown with the white square. i High magnification
COMP image of Type 3 FMA. Note that constituent minerals are much coarser than those of Types 1 and 2 FMAs. j High magnification COMP imageof chromian spinel grain shown in (d). k Magnified view of Ca and Al distribution in and around calcite vein in the chromian spinel grain shownin (j). Al-spl, Aluminous spinel; Cal, Calcite; Cpx, Clinopyroxene; Cr-spl, Chromian spinel; FMA, Fine-grained mineral aggregate; Native Fe, Native iron;
Ne, Nepheline; Ol, Olivine; Opx, Orthopyroxene; Pl, PlagioclaseAkizawa et al. Progress in Earth and Planetary Science(2024) 11:38(a) Na(b) K

Type 2Type 2

Page 7 of 28(c) CaType 2

Vein
(e)

(e)

Type 1TypeType 1

Type 1
Ty

pe

3

Ty

pe

3

1 mm

(d)

(e)

1 mm

1 mm

(e)

(f) Type 1

Vein

Ne

((g)
g)
(j)

Al-spl

2
pe
Ty

((f)
f)

Native Fe

Cr-spl
(h)

pe

Ty

Pl
Native Fe

1

Ol
Ne

Red: Fe
Blue: Na1 mm

3100 µm(g) Type 2

5 µm

(h)
(i)

Type 2OlPl
Native Fe

Type 3Al-spl

5 µm

Cal

Cal

Pl
Ol

Cr-spl

5 µm

Fig. 3 (See legend on previous page.)

100 µm(k)(k)

100 µm+((j)j)
Al-spl3
Fe -f ree Cr-spl(i) Type 3

50 µm

Red: Ca
Blue: Al

Cr-splAkizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 8 of 28

spinel, and native iron (Fig. 3g). Type 2 is composed of
the same minerals with Type 1, showing shape-preferredorientation (Fig. 3g). Native iron and nepheline are veryrare or even absent near the vein and in the Na-richmarginal zone (Fig. 3e), where Type 2 is developed andMg and Fe contents are poorer and richer than Type 1,
respectively (Fig. 2g, h). The grain size of Type 2 is mostly
identical to that of Type 1 (Fig. 3f, g). The minerals sealing the veins are elongate parallel to the vein direction
and their size is larger (up to ~ 10 µm in length) than thecommon Type 2 minerals (Fig. 3g). The olivine tends tooccur in the center of the veins and is fringed by plagioclase (Fig. 3g). Native iron is present sporadically in theveins (Fig. 2g; Supplementary Fig. S5).
Type 3 located near the olivine nephelinite showsdiverse mineral modes and microstructures (Fig. 3h, i).
Constituent minerals are olivine, spinel, and plagioclase
as in the cases of Types 1 and 2, whereas the volume fraction of sodic plagioclase is greater than those of Types
1 and 2, resulting in high Na concentration in Type 3(Figs. 2d, 3a). No native iron and nepheline are present
in Type 3 (Fig. 3h). The grain sizes of olivine, plagioclase,and spinel are ~ 10 µm across greater than those of Types
1 and 2. The olivine shows elongate morphology, forming
shape-preferred orientation radiating away from the areaof Type 2 (Fig. 3h).
4.2 Spinels in FMA

Fig. 4 Sub-micrometer-scale microstructures of Type 1 FMA in AI70
lherzolite. a Annular dark-field scanning transmission electron
microscopy (ADF-STEM) image taken with FE-TEM. b Elemental
distribution map of overlaying filters of red = Mg, green = Al,and blue = Ca acquired by FE-TEM-EDS. The numbers in a represent
analytical area or spots for quantification (see Supplementary
Table S1). Na-pl, Sodic plagioclase; Al-spl, Aluminous spinel; Ca-pl,
Calcic plagioclase, Ol, Olivine(Fig. 4). Large (~ 0.5 µm in diameter) spinel grains tend
to be in contact with olivine. Shape-preferred orientation
can be recognized for olivine, calcic plagioclase, and spinel (Figs. 3f, 4). Native iron and nepheline are euhedral
to subhedral and are locally in direct contact with eachother (Fig. 3f ).
Type 2 is developed around ~ 5 µm-thick veins thatsharply cut Type 1, forming ~ 50 mm-thick Na-rich marginal zones along the veins (the orange lines in Fig. 3c;a magnified view in Fig. 3e, g). The veins are branchedand construct a basic framework of network-like Type 2(Fig. 3a–d). The veins are sealed with olivine, plagioclase,

Spinel shows various modes of occurrence as describedabove. Two of the most abundant and important spinel
types are spinel incisions in FMA grains, and the otheris one of the essential constituents of FMA with olivine
and plagioclase. The former is more than a few hundredof microns in size (Fig. 3d) and Cr-rich (Cr#: Cr/(Cr + Al)
atomic ratio =  ~ 0.2). The latter is smaller than ~ a few
microns (Fig. 3f, g, i) and is Al-rich (Cr# =  ~ 0.05). We
hereafter distinguish the two spinel types by calling theformer chromian spinel and the latter aluminous spinel.
Two of the larger FMA grains out of 11 FMA grains
have the chromian spinel in the center (Fig. 2a–c). The
largest FMA grain includes several grains of chromianspinel with individual size of 0.1 to 1 mm in diameter(Fig. 3d). The chromian spinel grains are sparsely isolated
in two dimensions, forming an intricate aggregate with
an irregular highly indented morphology (Fig. 3d). Eachgrain is rounded to subrounded in shape (Fig. 3d). The
same FMA grain in a cut section ~ 1 mm above the thinsection plane also includes the chromian spinel grainsin the center (Supplementary Fig. S1). The size of FMA
grain in the cut section is ca. 6.0 mm in diameter (Supplementary Fig. S1e), which is slightly larger than the

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

FMA grain in the thin section (ca. 5.7 mm in diameter;
Supplementary Fig. S1f ), indicating that the former iscloser to the grain center in three dimensions. The chromian spinel grains in the cut section are compact andmassive, contrasting to grains in the thin section forminga sparsely intricate aggregate.
The crystallographic orientations of chromian spinel in the thin section were measured with a scanningelectron microscope equipped with an electron-backscatter diffraction (SEM-EBSD) system (see Supplementary methods for analytical detail). The chromian spinel
grains consist of distinct three clusters in terms of crystallographic orientations (Clusters A to C in Fig. 5). The
misorientations between the clusters range from 30 to
60°: ~ 30° between Clusters A and B, 40 to 50° between
Clusters A and C, and 50 to 60° between Clusters B andC, whereas those within each cluster are smaller than 20°(Fig. 5; Supplementary Fig. S6), indicating that the chromian spinel grains of Cluster A formed a single crystal
with branching morphology in three dimensions. Thisis consistent with the three-dimensional morphology ofchromian spinel as inferred above.
The chromian spinel grains in the largest FMA havesealed fractures (Fig. 3c, d). Two types of filling materials are recognized: one is calcite (see the red arrowheads
facing to each other in Fig. 3c) as identified by Raman
spectroscopic analysis (Supplementary Fig. S7; see Supplementary methods for analytical detail), and the otheris Na-rich (~ 8 wt% N­ a2O) glass (see the yellow arrowheads facing to each other in Fig. 3c). The fractures sealedwith calcite (calcite vein, hereafter) are oriented at highangles against elongation of chromian spinel grains(Fig. 3c, d). The calcite veins do not extend into the host
FMA and ends with a short (< 10 µm) protrusion into thehost FMA (Fig. 3j, k).
High concentration of calcite veins (see the dottedwhite circle in Fig. 3c) occurs on the extension of thelong extending veins in the center of Type 2 FMA (seethe orange lines in Fig. 3c), suggesting that the calcite
veins in the chromian spinel formed simultaneouslywith the veins in Type 2 FMA. In contrast, the veinssealed with Na-rich glass occur only in the chromianspinel grain placed in Type 3 FMA (see the yellow
arrowheads facing to each other in Fig. 3c). The FMA
bordering the chromian spinel grains belongs mostly
to Type 2 FMA, having width up to 0.2 mm, where
native iron and nepheline are absent or very minor inabundance (Fig. 3j), and Mg and Fe contents are lowerand higher than those of Type 1 FMA, respectively(Fig. 2g, h). The chromian spinel is locally in directcontact with Type 1 FMA (Fig. 3a–c), where the nativeiron and nepheline are as abundant as those of normal
Type 1 FMA.

Page 9 of 28

Fig. 5 Pole figure (PF) and inverse pole figure (IPF) map of chromianspinel in largest FMA of AI70 lherzolite. a PF of normal directions
to the {001} crystal family of planes of chromian spinel indexed in (b).
The color contrast of measurement points corresponds to the legend
color-coded according to the normal direction of the crystal plane
parallel to the X-axis. The reference frame of the X-, Y- and Z-axes
are the same as that in (b). The axes X and Y represent long axis
and short axis of the thin section, respectively. Equal-area, and lower
hemisphere projections. b IPF map of chromian spinel. The color
contrast of the measurement points corresponds to that in (a).
The analyzed grains can be grouped into three clusters in termsof similarity in crystallographic orientation (reddish Cluster A, paleorangish Cluster B, and pinkish Cluster C). To visually recognize
the similarity in crystal orientation relationship between Clusters
A to C, the data are rotated so that each of the normal axesof the (001) planes in Cluster A comes to each of the X-, Y-, and Z-axesof the pole figure. Original data before the rotation are presented
in Supplementary Fig. S6

5 Whole‑rock and Mineral Chemistry
5.1 Whole‑rock major elements, HSE, and Re–Os isotope
compositions

We conducted whole-rock chemical analyses of one olivine nephelinite (20190318A-11-27) and one lherzolite

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 10 of 28

Fig. 6 Bulk chemical compositions of AI70 lherzolite xenolith and its host volcanic rock from Aitutaki Island. a Mg/Si versus Al/Si plot (wt%). Dataof peridotite from fast-spreading mid-ocean ridges (MOR) including Garrett Transform Fault (Niu and Hékinian 1997), Hess Deep Rift (Regelous
et al. 2015), and Udintsev Fracture Zone (Niu 2004) are plotted. Peridotite data of cratonic region are from PetDB database (http://​www.​earth​chem.​org/​petdb), and those of Vitim volcanic field are from Litasov and Taniguchi (2002). Primitive upper mantle (PUM) data is from McDonoughand Sun (1995). b PUM-normalized bulk-rock highly siderophile element (HSE) patterns. Some of the patterns of Aitutaki and oceanic peridotites
that are similar to the pattern of AI70 lherzolite are highlighted by the thick pale blue and gray lines, respectively. PUM values are after Becker et al.
(2006). d Relationship between 187Os/188Os and ­Al2O3 content. Age reference data (Re-depletion model age: ­TRD) are shown with horizontal ticks
in the right-hand side. In (c) and (d), data of other Aitutaki samples are from Snortum et al. (2019) and those of oceanic peridotites are from Day
et al. (2017). d Total alkali versus silica (TAS) diagram. Data shown by blue circles are volcanic rocks from Aitutaki Island reported by Snortum et al.
(2019) for comparison. The thin lines represent rock classification by Le Bas et al. (1986) and Le Maitre (2002)(AI70). The AI70 lherzolite was powdered from a portion
one centimeter away from the host olivine nephelinite.
The analytical methods are summarized in Supplementary methods.
The whole-rock ­Al2O3 and CaO contents of AI70 lherzolite are 4.19 and 3.19 wt%, respectively (Supplementary
Table S2), and the Mg/Si and Al/Si weight ratios are 1.04
and 0.1, respectively, which are close to the PUM values(Fig. 6a). Reported whole-rock Mg/Si and Al/Si rations ofperidotite xenoliths from Aitutaki Island (Snortum et al.
2019) show a negative correlation and respectively range
from 0.9 to 1.3 and 0.04 to 0.12 (Fig. 6a), which overlap
with the ratios of peridotite xenoliths from continental

regions. The Mg/Si and Al/Si rations of AI70 lherzolite
indicate that it is one of the most fertile Aitutaki peridotite xenoliths (Fig. 6a).
The PUM-normalized HSE pattern of AI70 lherzolite ischaracterized by a smooth weakly upward convex pattern(Fig. 6b), which is comparable to some of the patterns
reported from Aitutaki peridotite xenoliths (Snortum
et al. 2019) and oceanic peridotites (Paquet et al. 2022)(see the thick pale blue and gray lines in Fig. 6b). The
187
Os/188Os ratio of AI70 lherzolite is 0.1248, whichcorresponds to 327 Ma T
­ RD (Re-depletion model age)
(Fig. 6c; Supplementary Table S2). Aitutaki peridotite xenoliths including AI70 lherzolite have diverse

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

187

Page 11 of 28

Os/188Os ratios ranging from 0.1184 to 0.1317 (Snortum et al. 2019) mostly overlapping with the variation ofoceanic peridotites (Fig. 6c).
The host olivine nephelinite is high in alkali contents(~ 5 wt% N­ a2O + ­K2O; Fig. 6d; Supplementary Table S2)
and classified as foidite in the total alkali versus silica(TAS) classification diagram according to Le Bas et al.
(1986). The major element composition is comparable tothat of the olivine nephelinite reported by Snortum et al.
(2019). The primitive upper mantle (PUM, a hypotheticalundifferentiated fertile upper-mantle reservoir)-normalized HSE pattern of olivine nephelinite is characterized
by a significant depletion of Os to Pd with slight depletion of Re relative to the PUM (Fig. 6b). The 187Os/188Os
ratio of the olivine nephelinite is 0.1615, which is muchhigher than the AI70 lherzolite and other Aitutaki mantle
xenoliths (Snortum et al. 2019) (Fig. 6c; Supplementary
Table S2).

grains occurring within the 1 cm-wide zone from thecontact with the olivine nephelinite (Figs. 2, 7). Suchspatial elemental variations of clinopyroxenes and grainboundaries sealed with glass or minerals are strongly
controlled by the distance from the contact with the olivine nephelinite. The chemical variations are characterized by abrupt changes at 1 cm from the contact withthe olivine nephelinite, indicating that the inner part of
AI70 lherzolite preserves pristine chemical compositionand has record of processes before the AI70 lherzolite
was entrapped in the olivine nephelinite magma, which
modified only the marginal zone of xenolith throughchemical diffusion and melt infiltration. In the following
subsection, we further examine grain-scale chemical zoning patterns of pyroxene grains in the inner part of the
AI70 lherzolite xenolith.

5.2 Mineral major‑element and trace‑element
compositions

We selected coarse orthopyroxene and clinopyroxene
grains away from the contact with the olivine nephelinite for examination of grain-scale heterogeneity. The distributions of Al in the orthopyroxene and clinopyroxene
grains are concentric and their Al contents increase fromthe core to the rim (Fig. 8a, c). The distribution of Ca in
the orthopyroxene grain is also concentric and its content
increases from the core to the rim (Fig. 8b), whereas theintragrain distribution of Ca in the clinopyroxene grain is
unclear (Fig. 8d).
We obtained line profiles of Al and Ca for the orthopyroxene and clinopyroxene grains (Fig. 8e, f ). The orthopyroxene grain is characterized by gradual increase in Al,
Ca, and Cr toward the rim (0.198 to 0.203 apfu = atoms
per formula unit ­AlO = 6, 0.037 to 0.039 apfu ­CaO = 6, and0.014 to 0.015 apfu C­ rO = 6) (Fig. 8e). The clinopyroxene issimilarly characterized by gradual increase in Al and Cr
toward the rim (0.275 to 0.282 apfu ­AlO = 6, and 0.027 to
0.028 apfu ­CrO = 6), whereas its Ca content shows monotonous increase from a rim to the opposite rim (Fig. 8f ).
Average core and rim compositions in the intervals indicated by the gray bands in Fig. 8e and f are given in Supplementary Table S3.

Spot and line analyses of major elements were conductedusing an EPMA, and mineral trace-element analysis was
conducted using a laser ablation-inductively coupledplasma-mass spectrometer (LA-ICP-MS). For LA-ICPMS analyses, an analytical area size of 60 µm was adopted
for clinopyroxene and FMA, and that of 120 µm wasadopted for orthopyroxene. The analytical methods are
summarized in Supplementary methods, and the results
are summarized in Supplementary Tables S3–S9.
5.2.1 Thin section‑scale elemental variation

AI70 lherzolite shows homogeneous distribution of elements over the thin section scale, excepting the contact
area with the host olivine nephelinite, which is obviously
noted by Na, Ca, and K maps (Fig. 2d–f ) as well as Al,
Mg, Fe, and Mg/Fe maps (Fig. 7). The Na and K heterogeneities are manifested by the variations of the grainboundaries sealed with glass or minerals. Sodium content along the grain boundaries sealed with glass is highin 1 cm-wide zone from the contact with the olivine
nephelinite, and sharply drops towards the inner zone,
where the grain boundaries are sealed with minerals(Fig. 2d). In contrast, K content along the grain boundaries sealed with glass show an opposite tendency, where
K content along grain boundaries is low in the 1 cm-widezone from the contact with the olivine nephelinite andabruptly increases toward the inner zone (Fig. 2e). The
high K content is attributed to the presence of K-feldspar
along the grain boundaries sealed with minerals (Supplementary Fig. S2h-j). Similarly, the Al, Ca, and Mg/
Fe maps show that high concentration of Ca, high Mg/
Fe, and low concentration of Al in the clinopyroxene

5.2.2 Zoning of major‑element composition in pyroxene
grains

5.2.3 Mineral major‑element composition

Average forsterite and NiO contents of olivine in AI70
lherzolite are 88.9 and 0.36 wt%, which are lower thanthose of the other peridotites from Aitutaki Island
reported by Snortum et al. (2019) (Fig. 9a; Supplementary Table S4). Olivine in the other modes of occurrences
was also measured for major elements. They are (1) finegrained olivine occurring in the decomposed orthopyroxene and clinopyroxene near the olivine nephelinite(Supplementary Fig. S2c–f ), (2) fine-grained olivine

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 12 of 28

Fig. 7 Pseudo-color distribution maps of AI70 lherzolite. a and b Pseudo-color maps of Al distribution with two-color coding schemes
for orthopyroxene and clinopyroxene, respectively. The positions of Fig. 8a–d are shown with the white squares. c and d Pseudo-color maps of Ca
distribution with two-color coding schemes for orthopyroxene and clinopyroxene, respectively. e and f Pseudo-color maps of Cr and Mg/Fe ratio
distributions, respectively. Warmer colors indicate higher elemental contents. The abbreviations are the same as Fig. 3

occurring along the grain boundaries sealed with minerals (Supplementary Fig. S2h–j), and (3) olivine phenocryst in the host olivine nephelinite. All these typesof olivine have lower forsterite and NiO contents thanthe coarse-grained olivine of the main lithology (Fig. 9a;
Supplementary Table S4). The average CaO content
and Mg# (Mg/(Mg + ­Fe2+) atomic ratio) of orthopyroxene are 1.0 wt% and 0.89, respectively (Fig. 9b; Supplementary Table S5). Average ­TiO2 and ­Na2O contents ofclinopyroxene are 0.68 wt% and 1.62 wt%, respectively(Fig. 9c; Supplementary Table S6). These values arebarely within the ranges of garnet lherzolites (Fig. 9b,c). The clinopyroxene of the main lithology is higher in­Na2O content than those of other modes in occurrences,which are (1) fine-grained clinopyroxene grains in thedecomposed orthopyroxene (Supplementary Fig. S2c,d), (2) spongy clinopyroxene near the host olivine
nephelinite (Supplementary Fig. S2e, f ), (3) fine-grainedAkizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 13 of 28

Fig. 8 Pseudo-color distribution maps and line profiles of orthopyroxene and clinopyroxene in AI70 lherzolite. a and b Local pseudo-color mapsof orthopyroxene for Al and Ca, respectively. c and d Local pseudo-color maps of clinopyroxene for Al and Ca, respectively. Positions are shownwith the white rectangles in Fig. 7a and b. Warmer colors indicate higher elemental contents than cooler colors. e and f Line profiles of Al, Ca,
and Cr in the selected orthopyroxene and clinopyroxene grains along the red lines in (a) to (d). Gray bands represent data used to calculate mean
compositions for core and rim listed in Supplementary Table S3. The abbreviations are the same as Fig. 3. The cation data are calculated basedon the assumption that the oxygen equals to 6

clinopyroxene along the grain boundaries sealed withminerals (Supplementary Fig. S2h–j), and (4) clinopyroxene in the groundmass of the host olivine nephelinite(Fig. 9c; Supplementary Table S6). The chromian spinelin the FMA is fringed by a few tens of micrometer-thickferric iron-free aureole (Figs. 3k, 9d). The calcite fillingthe fracture developed in the chromian spinel (Fig. 3j,k) and that occurring along the grain boundaries sealedwith minerals (Supplementary Fig. S2h–j) are poor in Mg(MgO < 0.5 wt%), although the former locally contains ~ 9
wt% MgO (Supplementary Table S8).
Bulk major-element composition of Types 1 and 2
FMAs were obtained by EPMA with defocused spot
analysis and processing map analysis (Table 1; See Supplementary methods for details of analyses). The two
methods give consistent results, which are similar tochemical compositions of garnets with pyrope component as high as 70% reported from garnet lherzolite

xenoliths (Table 1). Type 2 FMA contains higher Na
and lower K (~ 2.5 wt% ­Na2O, and 0.4 wt% ­K2O) thanType 1 FMA (~ 1.7 wt% ­Na2O, and 2.8 wt% ­K2O)(Table 1). We modeled the bulk FMA composition fromidentified mineral compositions by optimizing theirmodal abundances (see Supplementary Table S10). The
modeling results show that the bulk chemical compositions of Types 1 and 2 FMAs can be reproduced by theidentified minerals with their volume fractions nearlyconsistent with the observed ones (Supplementary
Table S10).
5.2.4 Mineral trace‑element composition

Chondrite-normalized rare-earth element (REE) patterns of large orthopyroxene grains are marked bygently sloped segment from middle REE (MREE) toheavy REE (HREE) and clear depletion from MREE tolight REE (LREE) irrespective of core and rim (Fig. 9e).

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 14 of 28

Fig. 9 Chemical compositions of minerals in AI70 lherzolite. a NiO content of olivine plotted against forsterite content [Fo = 100 Mg/(Mg + ­Fe2+)
atomic ratio]. b CaO content of orthopyroxene plotted against Mg# [Mg/(Mg + ­Fe2+) atomic ratio]. c ­Na2O content of clinopyroxene plottedagainst ­TiO2 content. d Cr-Al-Fe3+ triangular diagram for spinel. e CI chondrite-normalized rare-earth element (REE) patterns of orthopyroxene
and clinopyroxene. f CI chondrite-normalized REE patterns of FMA. Red squares in (a) to (d) are for core compositions of the main constituent
minerals of the main lithology. Open squares in (a) and (c) are for fine-grained olivine and clinopyroxene grains occurring in decomposed
orthopyroxene, respectively (Supplementary Fig. S2c, d). Small solid black squares in (a) and (c) are for respectively fine-grained olivine
and clinopyroxene grains occurring in decomposed spongy clinopyroxene, respectively (Supplementary Fig. S2e, f ). Open red triangles in (a) and (c)are for olivine and clinopyroxene grains occurring along grain boundaries sealed with minerals (GBM) coexisting with calcite (Supplementary
Fig. S2h–j). Small solid red triangles in (a) and (c) represent core compositions of olivine phenocryst and groundmass clinopyroxene in the hostolivine nephelinite, respectively (plotted outside of the diagrams as indicated by red arrows with values). Open square and small solid squarein (d) are for rim composition of chromian spinel and fine-grained aluminous spinel in FMA, respectively. Error bars are one standard deviation,some of which are smaller than the symbols. In (a), reported olivine compositions of Aitutaki lherzolites and harzburgites are plotted by the open
blue circles (Snortum et al. 2019). Mineral data of peridotite xenoliths from Vitim volcanic field are referred from Litasov and Taniguchi (2002)
for comparison. Garnet compositions and coexisting clinopyroxene compositions of Vitim garnet lherzolites are shown in (e) and (f) for comparison
with those of FMA and clinopyroxene in AI70 lherzolite. CI chondrite values in e and f are from McDonough and Sun (1995). The abbreviations are
the same as Fig. 3

Chondrite-normalized REE patterns of large clinopyroxene grains are characterized by steady enrichment from HREE to LREE irrespective of core andrim (Fig. 9e). The spongy clinopyroxene has a steeper
REE pattern with higher contents than the large clinopyroxene grains and is characterized by a positive Eu
anomaly (Fig. 9e). The chondrite-normalized REE patterns of bulk Type 1 FMA decline steeply from HREE to
LREE, which is comparable to those of garnet in garnet
lherzolite with slightly higher values of LREEs (Fig. 9f;
e.g. Litasov and Taniguchi 2002). Type 2 FMA has thesame MREE to HREE contents, but is richer in LREE
than Type 1 FMA, showing spoon-shaped REE patterns(Fig. 9f ).

6 Garnet origin of FMA
Snortum et al. (2019) reported “nearly opaque aggregates of very fine-grained minerals” that exhibit amorphous or grainy texture at micron scale from two fertile
Aitutaki lherzolite xenoliths in olivine nephelinites (samples AK1023A and AK1023B). The aggregates enclose
chromian spinel grains and have sporadic grains of
Fe–Ni metals. The petrographic characteristics of the“nearly opaque aggregates of very fine-grained minerals” are essentially the same as those of FMA describedin this study [cf. Figure 2a, b of Snortum et al. (2019) and
Fig. 2a, b herein]. It is presumed that the “nearly opaqueaggregate of very fine-grained minerals” is an FMA. The
similarities of whole-rock chemical compositions andAkizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 15 of 28

Table 1 Bulk major-element composition of Type 1 FMA and Type 2 FMA in comparison with garnet (wt%)
Sample name

AI70

YL-2

V-860

Locality

Aitutaki, Cook Islands

Vitim volcanic
field

Note

Type 1 FMA

Analytical number

4

Area 1

93

Area 2Analytical technique

Defocused spot

Chemical mapping technique(See Supplementary Fig. S5)

Defocused spot

Chemical mapping technique(See Supplementary Fig. S5)

GarnetType 2 FMA

SiO2

42.8

42.6

42.9

42.6

40.7

42.3

TiO2

0.22

0.22

0.12

0.23

0.16

0.45

Al2O3

22.7

22.9

22.9

23.0

21.7

22.9

Cr2O3

1.41

1.22

1.18

1.26

0.94

0.57

FeO*

4.46

6.10

4.82

6.63

6.70

9.01

MnO

0.37

0.29

0.33

0.26

0.18

0.26

MgO

21.1

20.3

19.7

19.1

20.0

20.3

CaO

4.4

4.8

6.0

5.4

5.37

4.32

Na2O

1.28

1.73

2.47

2.41

2.22

K2O

2.77

Not analyzed

0.36

Not analyzed

Total

101.5

100.2

100.7

100.9

98.0

100.1

Mg#

0.894

0.842

0.879

0.856

0.842

0.801

Bulk major-element compositions of Type 1 FMA and Type 2 FMA were determined either by (1) defocused spot-analysis (20 µm in diameter) with EPMA, or (2)chemical mapping technique using QntMap developed by Yasumoto et al. (2018). Analyzed areas (Areas 1 and 2) are indicated in Supplementary Fig. S5. See
Supplementary methods for analytical details. FeO* is total iron as FeO. Garnet compositions of Vitim garnet lherzolite xenoliths are referred from Litasov andTaniguchi (2002)

petrographic features, such as grain size, strong foliation
or lineation, and the presence of chromian spinel onlyin the FMA indicate that AI70 lherzolite and AK1023A
and AK1023B lherzolites are fragments derived fromthe same mantle material. Snortum et al. (2019) inferredthat the FMA was “orthopyroxene” and interpreted
that the “orthopyroxene” and enclosed chromian spinel
grains were formed by a reaction between garnet andolivine by decompression from the garnet-stability field
to the spinel-stability field. However, our petrographic
and mineral chemical data of the FMA indicate that the
FMA was not orthopyroxene. The bulk chemical composition of FMA is similar to pyrope-rich garnet in thegarnet peridotites from the cratonic mantle with regard
to the major elements (Table 1) and REEs (Fig. 9f, Supplementary Table S9). The strongly LREE-enriched andHREE-depleted pattern of clinopyroxene evidently suggests its coexistence with garnet (Fig. 9e). Snortum et al.
(2019) described that their FMA has a bulk composition
approaching to orthopyroxene, but their reported bulk
composition analyzed with EDS is very different fromthat of orthopyroxene and rather similar to pyrope-richgarnet as it is the case for our FMA (Table 1). Therefore,
we conclude that the FMA was pyrope-rich garnet before
decomposition into fine-grained mineral aggregate.

The grain size of Type 1 FMA-forming minerals(< ~ 2 µm) (Figs. 3f, 4) is much smaller by three ordersof magnitude than the chromian spinel inclusions in the
FMA, which are ~ 1 mm across (Fig. 3d). From this fact
combined with our contention mentioned above, we
argue that the FMA is a breakdown product after unstable pyrope-rich garnet during or shortly before entrapment as xenolith in ascending olivine nephelinite magma
and that the chromian spinel was enclosed in the garnet
before its breakdown into fine-grained mineral aggregate with maintaining garnet composition (cf., Obataet al. 2013, 2014). Such isochemical breakdown of garnet
is supported by our least-squares mass balance calculation (see Supplementary Table S11). The modeling results
show that the bulk chemical compositions of Types 1 and2 FMAs can be reproduced by components of garnet if­Na2O and ­K2O contents are ignored or minor. It is presumed that Na and K were introduced during or after
the breakdown of garnet as discussed in Sect. 8.5. Thus,
we conclude that all mineral phases of the major lithology such as pyroxenes in the AI70 lherzolite had been in
equilibrium with the garnet shortly before the extraction
as xenolith. We present geothermobarometry of AI70
lherzolite in Sect. 7 under the assumption that the AI70
lherzolite resided in the garnet-stability field.

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

7 Geothermobarometry: pressure
and temperature shortly before rapid xenolith
exhumation
Chemical zonings of pyroxenes are caused by changes
in pressure, temperature, and/or bulk chemical composition. Such changes in environment affect mineralchemical compositions of the outermost rim and aretransmitted inside the grain via chemical diffusion, causing a concentric zoning pattern unless recrystallization or
grain boundary migration (dissolution or growth) due todeformation and transfer reactions with a grain-boundary fluid phase occur. We investigated chemical zoningsof large pyroxene grains from the center of AI70 lherzolite (Figs. 7, 8), where mineral grains are least affectedby chemical modifications during events of xenolith formation and extraction such as fragmentation of mantle
materials and their entrapment into a magma followed bytransportation to the surface. The selected large orthopyroxene grain shows nearly concentric zoning of Al, Ca,
and Cr increasing from the core to the rim (Fig. 8e). The
selected large clinopyroxene grain shows similar concentric zoning of Al and Cr increasing from the core to
the rim, though Ca is homogeneous in composition withminor asymmetry with maximum difference in CaO
contents at both rims by ca. 0.007 apfu C­ aO=6, which is
negligibly small less than 1% relative to the mean value of0.720 apfu C­ aO=6 (Fig. 8f; Supplementary Table S3). The
weak asymmetry of Ca zoning may be due to localized
Ca heterogeneity along grain boundaries and does notreflect pressure and temperature changes that the entire
rock underwent.
The orthopyroxene grain is zoned in Ca over the grain
scale (~ 2 mm in diameter), which cannot be modified
during xenolith extraction process if Ca diffusivity inorthopyroxene (Cherniak and Liang 2022) and short timescale of xenolith transportation (Spera 1984) are taken
into consideration. Moreover, the concentric zoning of
Al and Cr in both pyroxenes cannot be modified during
the xenolith extraction processes, as the diffusivities of Al
and Cr are much smaller than that of Ca (Cherniak andDimanov 2010). The rim compositions of concentricallyzoned pyroxene grains, therefore, record pressure andtemperature (P–T) conditions shortly before the xenolith was entrapped in a magma for transportation to thesurface.
The concentric zoning patterns of Ca and Al in thepyroxenes (Fig. 8) suggest decompression and/or heating before xenolith extraction since the AI70 lherzolite was in the garnet-stability field before its extraction(Gasparik 2003). The Ca, Al, and Cr profiles of theselected large orthopyroxene grain do not have a nearlyflat core region in any of these elements (see Fig. 8a for
Al and Fig. 8b for Ca), indicating that an application ofPage 16 of 28

geothermobarometry to the core does not provide ameaningful P–T condition for long residence. The Al
and Cr profiles of the selected large clinopyroxene grain
also do not have a wide flat core region, which is consistent with those of the orthopyroxene if smaller grain sizeand slower Al and Cr diffusivity of clinopyroxene thanorthopyroxene are considered (Fig. 8). Hence, the zoning
patterns of pyroxenes (Fig. 8e, f ) suggest that a plausible
initial P–T condition is higher in pressure and/or lower
in temperature than that calculated from the pyroxene
cores. In order to extract P–T information from the coresof pyroxenes, initial P–T condition must be constrainedby reproducing the observed zoning profiles with a diffusion model, which will be presented in Sect. 8.1.
We used two-pyroxene thermometry formulated by
Brey and Köhler (1990) ­(TBKN) because it is best calibrated for natural mantle peridotites and often used in
the literatures. As discussed above in Sect. 6, the FMA
was pyrope-rich garnet, and thus the pressure dependence of Al solubilities in the pyroxenes should be evaluated under the presence of garnet. We employed
Al-in-orthopyroxene barometer formulated by Brey andKöhler (1990) ­(PBKN) to estimate pressure. The chemicalcompositions used for our P–T estimate are mean values
listed in Supplementary Table S3.
We applied the thermobarometry to the rim-rim pair ofpyroxenes and obtain 1140 ± 3 °C and 2.12 ± 0.02 GPa byusing the rim of low-Ca side of clinopyroxene. We calculated P–T for the core-core pair knowing that the application of geothermobarometry to the core-core pair does
not provide a meaningful P–T condition for long residence as explained above and obtained 1138 ± 3 °C and2.14 ± 0.03 GPa. The uncertainty of P–T estimates (± 1σ,
68% confidence interval) is evaluated with a Monte-Carlo
method assuming that the source of errors is derived
from analytical uncertainties with Gaussian distributions. The results are the same irrespective of which rim
compositions of clinopyroxene are used. The estimated
P–T conditions are within the garnet-stability field (Gasparik 1987; Ziberna et al. 2013), being consistent withthe assumption that the AI70 lherzolite was derived fromthe garnet-stability field as contended above. The Cr#of the chromian spinel is as high as 0.2 (Supplementary
Table S7), which is slightly higher than that of spinel in
fertile spinel lherzolite (Arai 1987) and is expected if spinel coexists with garnet in the spinel-garnet transition
conditions (Nickel 1986). Although the two P–T estimates are within the errors, the Al and Ca zoning profiles
without a flat core region particularly for Ca (Fig. 8a, b, e)
and the contrasting P–T values calculated from the corecore and rim-rim pairs indicate that the core-core pairrecorded higher pressure and lower temperature thanthat registered by the rim-rim pair.

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

8 Discussion
8.1 Pressure and temperature history before extractionof AI70 lherzolite xenolith

Gradational chemical zoning patterns of Al, Ca, and Cr
recognized in the coarse orthopyroxene and clinopyroxene grains in AI70 lherzolite (Fig. 8) require decompression and/or heating in the garnet-stability field asmentioned in Sect. 7. To quantitatively estimate the P–T
history of AI70 lherzolite before xenolith extraction, weused a diffusion model and tried to constrain P–T pathand initial P–T condition under a constant decompression rate by reproducing the observed Al and Ca zoning
patterns in the large orthopyroxene grain (Ozawa 1997,
2004; Aoki et al. 2020). Because the rim compositions
reflect the P–T condition shortly before the xenolith
extraction, the P–T path was constrained to reach theP–T condition estimated above for the rim-rim pair.
Successful modeling results are presented in Fig. 10. We
examined various P–T paths, initial P–T conditions, anddecompression rates. The decompression rate was varied
from 8 × ­10−8 to 6 × ­10−7 GPa ­year−1. Based on numerous
simulation runs, we came to conclude that the observed
zoning patterns can only be reproduced by decompression on the time scale of a few tenths of million years(0.33–0.81 million years) at decompression rates ranging
from 8 × ­10−8 to 2 × ­10−7 GPa ­year−1, which corresponds

Page 17 of 28

to exhumation rate from 0.24 to 0.6 cm ­year−1) (Fig. 10b).
The discrepancies between the observed zoning profiles
and modeled profiles might be due to a change in decompression rate during the ascent. We found that a heatingin the later stage following a short isothermal decompression (Fig. 10b) is required to account for the increase in
Ca content of orthopyroxene towards the rim (Fig. 10a).
The initial pressure condition can be higher than 2.16
GPa as assumed from the absence of flat core regions in
the zoning profiles (Fig. 8e) and modeling results shownin Fig. 10 when a slower ascent rate to reach the final
depth is assumed. We conclude that the AI70 lherzolite was weakly heated during slow ascent (ca. less than
2 × ­10−7 GPa ­year−1) shortly before the rapid extraction
as xenolith by a magma. The heating during decompression is consistent with the differences in P–T values calculated from the rim-rim and core-core pairs. Since theP–T dependence of Cr solubility is similar to Al in thepyroxenes (Aoki et al. 2020), the zoning patterns of Cr
may be explained by the same P–T history.
8.2 Origin of chromian spinel and pyrope‑rich garnet

We demonstrated above that the FMA was pyrope-richgarnet and the AI70 lherzolite resided in the garnetstability field before it was rapidly extracted as xenolith.
The chromian spinel is always enclosed in pyrope-richFig. 10 Modeled Ca and Al zoning in 2.1 mm-sized orthopyroxene grain in AI70 lherzolite. Zoning profiles are shown in Fig. 8e. The analyzed Ca
and Al contents are normalized by the core compositions and are plotted as black dots. The initially homogeneous orthopyroxene is assumed
to have started decompression at pressure higher than the calculated pressure for the core-core pyroxene pair, whose A­ l2O3 is higher
than the modeled core value. Modeling results of adiabatic upwelling with weak heating shortly before xenolith entrainment are shown in (a)and (b). The xenolith was entrapped in a magma at the depth shown by the black diamond in (b). Three different constant decompression rates
of 8 × ­10−8, 2 × ­10−7, and 6 × ­10−7 GPa ­year−1 were examined in the modeling. Their pressure changes with time are shown by the colored paths
denoted as “P-time”, and those with temperature are shown by the black paths denoted as “P–T” in (b). The zoning patterns of orthopyroxene are
best reproduced by adiabatic mantle upwelling with weak heating at slow decompression rates (8 × ­10−8 to 2 × ­10−7 GPa y­ ear−1 corresponding
to 0.6 to 0.24 cm year.−1) at least in the latest stage of decompression. Isopleths of Ca and Al contents per 6 oxygens in orthopyroxene arerespectively shown with the solid lines and dotted lines in the P–T space of (b) (Gasparik 2000; Aoki et al. 2020)

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

garnet (Fig. 2c). This garnet-spinel relationship suggeststhat the chromian spinel was a relict from incomplete
consumption of spinel during pyrope-rich garnet formation induced by compression from the spinel-stability field to the garnet-stability field. The relict origin of
the chromian spinel is supported by two petrographic
observations.
First, the rounded to subrounded chromian spinel
grains isolated in two dimensions in the largest FMA
have a similar crystallographic orientation to form thelargest cluster (Cluster A; Fig. 5), indicating that thecluster is a single crystal with very irregular and intricate three-dimensional morphology. The same FMA ina section ~ 1 mm above the plain of the thin section has
slightly smaller, more compact, and massive morphology
than those in the thin section (Supplementary Fig. S1e,f ). Such morphological difference suggests that the chromian spinel has a compact and massive core decorated byan irregular highly indented fringe with crystallographic
continuity in three dimensions. The morphology can be
explained either by rapid growth or dissolution, but thelatter possibility is more plausible because of the presence of the compact and massive core region and theoccurrence completely isolated within the FMA. Second,the FMA is surrounded by olivine-rich and pyroxenepoor aureole with a few millimeters in thickness (Fig. 2a,b; Supplementary Fig. S1d). The chromian spinel grains
with irregular highly indented fringe in the pyrope-richgarnet surrounded by olivine-rich zone can be attributed
to the consumption of spinel and pyroxenes with the formation of garnet and olivine. The reaction is consistent
with the following pressure-sensitive reaction (e.g., Kushiro and Yoder 1966):
Spinel + Orthopyroxene + Clinopyroxene = Garnet + Olivine.

(1)
The low-pressure stability limit of garnet peridotite for
the CaO–FeO–MgO–Al2O3–SiO2 (CFMAS) system andCr-bearing natural system (MORB Pyrolite) are shownby the solid pink lines in Fig. 11a. Reaction (1) proceeds toward the right-hand side either by temperature
decrease or pressure increase (Fig. 11a). If the AI70 lherzolite was derived from a domain within the lithosphericmantle, the formation of garnet can be explained by asimple cooling of the oceanic lithosphere (e.g., Neal andNixon 1985). By contrast, a heat loss and thus temperature drop of a domain in the asthenospheric mantle is
implausible at a constant depth because there is no heatsink in the asthenosphere. However, if a domain in theasthenospheric mantle adiabatically descended by convective movement, a pressure increase with trivial temperature increase may have driven the reaction (1) to theright.

Page 18 of 28

As discussed in Sect. 7, the concentric zoning patternsof Ca and Al in the large pyroxene grains in the AI70 lherzolite (Fig. 8) can be reproduced by a pressure decrease in
the garnet-stability field before quick xenolith extraction(Fig. 10b). The absence of flat core regions even in the Al
and Cr zoning profiles in the pyroxenes substantiates thatthe pressure decrease continued from a deeper level thanthat estimated from the core-core pyroxene pair (the blue
arrow pointing the black diamond 3 in Fig. 11a). It is, therefore, inferred that the mantle domain from which the AI70
lherzolite was extracted underwent a decompression fromthe depth deeper than ~ 70 km (2.16 GPa in pressure).
There is no plausible mechanism to produce such uplift in
the cooling oceanic lithosphere, negating the lithosphereorigin of AI70 lherzolite. Therefore, we conclude that the
AI70 lherzolite was derived from a domain in convecting
asthenosphere adiabatically ascended just before the xenolith extraction.
Based on the above discussions, we further envisagethat an asthenospheric mantle having large equant aluminous spinel grains coexisting with pyroxenes and olivine was dragged down from the spinel-stability field
to the garnet-stability field deeper in the mantle. The
downward movement drove the reaction (1) toward theright, resulting in the transition of aluminous spinel into
pyrope-rich garnet from the margin by decomposing surrounding pyroxenes and forming olivine (see the panels
1 and 2 in Fig. 11b). Further advance of the reaction (1)probably slowed down and eventually stopped to leave
chromian spinel inclusions with irregular highly indented
shape as a relict in the grown garnet (see the panel 2 inFig. 11b). The slowing and ultimate cessation of the spinel consumption was probably due to the increase in
Cr# of the relict chromian spinel from ~ 0.1, mean valueof fertile spinel peridotites (Arai 1987), to ~ 0.2 (Supplementary Table S7) and the inhibition of reaction progressby complete isolation of spinel from the outside of thegarnet grain. The microstructure of the chromian spinelcored garnet requires preferential formation of garnet on
spinel surface (see the panel 2 in Fig. 11b). Similar microstructural relationships between spinel and garnet havebeen observed in a variety of different settings: garnetfringing spinel in websterite from the Nikubuchi peridotite body in the Sanbagawa metamorphic belt, NW Japan(Yokoyama 1980); garnet-fringing spinel in a pyroxenite
xenolith from the Middle Atlas, Morocco (Moukadiri andKornprobst 1984); and corroded spinel relict in garnet
in mylonitic garnet harzburgite from the Cabo Ortegalultramafic complex, NW Spain (Ibarguchi et al. 1999).
8.3 Formation process of type 1 FMA

The FMA is composed mostly of fine-grained olivine,
calcic and sodic plagioclases, and aluminous spinelAkizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 19 of 28

Fig. 11 Inferred thermal and decompression history of AI70 lherzolite. a Inferred pressure–temperature path of xenolith exhumation. Last P–T
condition shortly before xenolith extraction is marked by the black diamond 3 at 1140 °C and 2.12 GPa. A path of adiabatic mantle upwelling
inferred from our diffusion modeling (see Fig. 10) is portraited with the pale blue arrow. The black diamonds 1 to 7 correspond to those in (b).
See Sect. 8.7 for further details of the inferred P–T history that the AI70 lherzolite passed through in the convecting asthenospheric mantle. The
mineral phase transition from garnet peridotite to spinel peridotite (grt-spl) for the Ca-Fe–Mg-Al-Si (CFMAS) system with Mg/(Mg + Fe) = 0.9
after Gasparik (1987) and that for MORB Pyrolite system after Ziberna et al. (2013) are shown with the pink lines. The mineral phase transitions
from spinel peridotite to plagioclase peridotite for the Na-Ca-Mg–Al-Si (NCMAS) system with anorthite contents (100 × Ca/(Ca + Na) atomic
ratio) of plagioclase = 80, 50, and 30 after Gasparik (1987) are shown with the pale blue lines. The solidus determined by melting experimentsof volatile-free peridotites and volatile-present peridotites are shown with the thin black lines (Falloon and Green 1989; Hirschmann 2000; Dasgupta
and Hirschmann 2006). The thin broken black lines represent adiabatic temperature gradients for solid mantle after Herzberg and Asimow(2015). The lithosphere and asthenosphere boundaries inferred for the Southwest Pacific at ~ 120 millions of years (Herath et al. 2022) and for the
Northwest Pacific at ~ 130 millions of years (Kawakatsu et al. 2009) are presented by the brown lines. b Inferred mineral phase transitionand decomposition history to form Types 1 to 3 FMAs after pyrope-rich garnet in AI70 lherzolite. See Sects. 8.3 and 8.4 for further details. Al-spl,
Aluminous spinel; Cal, Calcite; Cpx, Clinopyroxene; Cr-spl, Chromian spinel; FMA, Fine-grained mineral aggregate; Grt, Garnet; Native Fe, Native iron;Opx, Orthopyroxene; Pl, Plagioclase; Spl, Spinel(Fig. 4), which were transformed from pyrope-rich garnet. The upper pressure limit of plagioclase peridotite in
the CaO–MgO–Al2O3–SiO2 (CMAS) and ­Na2O-CMAS(NCMAS) systems are shown by the solid pale blue lines
in pressure–temperature space (Fig. 11a). The garnet-likebulk composition of Type 1 FMA can be reproduced bygarnet components if alkalis are excluded (Supplementary Table S11), which implies isochemical breakdown ofgarnet and later introduction of alkalis. It is known thatthe garnet is isochemically decomposed into symplectic mineral aggregate by the following reaction duringdecompression (Obata et al. 2013):

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Garnet → Orthopyroxene + Plagioclase + Spinel.

(2)
Reaction (2) can start in the spinel-stability field
above the lower pressure limit of garnet stability, forming orthopyroxene + plagioclase + spinel assemblage.
The coexistence of olivine and calcic plagioclase (anorthite content up to 87.4; Supplementary Table S1) in
Type 1 FMA indicates that the isochemical decomposition of garnet in the spinel-stability field was followed
by transformation of orthopyroxene-bearing assemblage
into the orthopyroxene-absent/olivine-bearing assemblage. The isochemical garnet breakdown to produce
the FMA should have continued at pressures lower than
1.0 GPa, where plagioclase + olivine assemblage is stable(Fig. 11a). The pressure where the AI70 lherzolite resided
shortly before the quick xenolith extraction estimatedfrom the rim-rim pair of pyroxenes is 2.12 GPa. The presence of plagioclase + olivine assemblage with small grain
size of constituent minerals (~ 5 µm) in Type 1 FMA indicates that the AI70 lherzolite was quickly decompressedfrom ~ 2 GPa to below 1 GPa (Fig. 11a), attaining isochemical breakdown of garnet into the olivine + plagioclase + spinel assemblage observed as the FMA.
Experiments examining the decomposition of garnet showed that time scales to produce a polyphaseintergrowth of orthopyroxene + spinel + anorthite in
the size up to few microns are few days at temperaturesof ~ 1050 °C (Ezard et al. 2022). The time scale of the isochemical decomposition of garnet documented in thisstudy is inferred to be on the same order. Such shorttime scale of garnet decomposition is possible only bytransportation by a magma. We hence suggest that the
breakdown of garnet was driven by decompression during entrapment of the xenolith in a magma and its rapidtransportation to the surface.
The olivine nephelinite hosting the AI70 lherzolite
represents a magma that transported to the surface inthe final stage of quick xenolith extraction. Judging fromthe fact that the olivine nephelinite has Fe-rich olivine
phenocrysts (Fig. 9a, Supplementary Table S4), it was
not equilibrated with the mantle, but was fractionatedto some degrees from a parental magma. The parental
magma is expected to have originated in the mantle at apressure higher than ~ 2 GPa, where the AI70 lherzolite
resided before its quick transportation. Experimentally
determined liquidus temperature of olivine nephelinite
magma is ~ 1400 °C at 2.25 GPa (Bultitude and Green1971). By adopting viscosity of 2 Pa ­s−1 estimated foralkali basaltic magma at high pressure and temperature (1.4 GPa and 1440 °C, Bonechi et al. 2022) and thexenolith size of 3 cm in diameter, a velocity of magma
ascent must be faster than 0.1 m ­s−1 to bring the AI70

Page 20 of 28

lherzolite to the surface (e.g., Spera 1984; Russell andJones 2023). If the olivine nephelinite magma transported
the AI70 lherzolite from 70 km depth (Fig. 11a), timescale of transportation is shorter than ~ 8 days. This timescale is a maximum value as the parental magma of theolivine nephelinite that initiated the xenolith transportation could have had lower viscosity and density. Such
possibility is discussed in the following subsections after
clarifying the origins of Types 2 and 3 FMAs and grainboundaries sealed with glass and mineral aggregate.
8.4 Formation process of type 2 FMA and type 3 FMA

Chemical and microstructural relationships between
three types of FMA show that Type 2 FMA postdated
Type 1 FMA, and Type 3 FMA postdated Type 2 FMA.
Since Type 2 FMA develops from the FMA margin andalong the network-like fractures in the FMA grains(Fig. 3c, e), and Type 3 FMA occurs only in the FMA
margin facing the contact with the olivine nephelinite(Figs. 2i, 3c), the formation of Types 2 and 3 FMAs wasenhanced by infiltrations of melt or fluid. Consideringthat the chemical characteristics of Types 2 and 3 FMAs
are distinct as shown in Sect. 4, the fluids responsible forthe formation of the two types of FMA are different inchemical composition as well as in physical properties.
The localized occurrence of Type 3 FMA confinedin ~ 1 cm-wide zone from the contact with the olivine
nephelinite (Fig. 2i), and its latest formation than theother types of FMA implies that Type 3 FMA formed byan infiltration of the olivine nephelinite magma during
the quick xenolith extraction to the surface. This is substantiated by its coexistence with open grain boundaries
sealed with glass (Fig. 2d–i), which will be further discussed in Sect. 8.5.
By contrast, the fluid responsible for the formation ofType 2 FMA cannot be the olivine nephelinite magma.
This is because the chemical composition of Type 2 FMA
is distinct from that of Type 3 FMA and because its formation stage predated the formation of Type 3 FMA. The
margin of each FMA grain is occupied by Type 2 FMA
without exception (Fig. 2d–i), and the spatial variationsof Na and K along the contacts between Types 1 and 2
FMAs are gradational (Fig. 3a, b). Such petrographic andchemical characteristics suggest that a fluid pervasively
infiltrated into each FMA grain, replacing Type 1 FMA
from the grain margin and further along fractures toform a dendritic network of Type 2 FMA (Fig. 3a–c, e, g).
Because there exists the area of high concentration of calcite veins in the chromian spinel inclusions on the extension of network-like veins, from which Type 2 FMA isdeveloped (Fig. 3c), Type 2 FMA was formed in responseto the infiltration of a carbonaceous fluid with very low
viscosity and surface tension. The open grain boundaries

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

sealed with minerals may be related to the formation ofType 2 FMA, which will be further discussed in Sect. 8.5.
8.5 Origin of grain boundaries sealed with glassor minerals and involved fluid/melt

The AI70 lherzolite has open grain boundaries, which
are sealed with glass or minerals (Fig. 2d–i), suggestingthat infiltration of fluid or melt along grain boundaries
opened due to rapid decompression. The glass and mineral-sealed grain boundaries have intimate relationships
with Types 3 and 2 FMAs, respectively, in terms of spatial
distribution and chemical and mineralogical characteristics as described above. Therefore, the glass or mineralsealing the grain boundaries have unique information toreveal the nature of fluid/melt responsible for the formation of Types 2 and 3 FMAs.
The olivine nephelinite has high ­Na2O content up to 4.9
wt% (Fig. 6d; Supplementary Table S2). The distributionof grain boundaries sealed with Na-rich glass (Fig. 2d;
Supplementary Fig. S2g) coincides with the presence of
decomposed orthopyroxene and spongy clinopyroxene(Supplementary Fig. S2c–f ), and Type 3 FMA, in which
native iron is absent and the constituent mineral grains
are as coarse as up to ~ 10 µm (Fig. 3h, i). All these characteristics are confined within the 1 cm-wide zone fromthe contact with the olivine nephelinite (Figs. 2, 7), andhence indicate that the olivine nephelinite magma infiltrated along grain boundaries and induced decomposition of orthopyroxene and clinopyroxene with theformation of Type 3 FMA. The thickness of partiallydecomposed orthopyroxene and clinopyroxene is lessthan 1 mm (Supplementary Fig. S2c–f ), which can beformed on the time scale of ~ 10 days (Tsuchiyama 1986).
This time scale is comparable to that of xenolith transportation (a few days) as estimated in Sect. 8.3.
The grain boundaries sealed with minerals are locatedinside the zone of the grain boundaries sealed with Narich glass (Fig. 2; Supplementary Fig. S2g). The minerals filling the grain boundaries are calcite, olivine,clinopyroxene, K-feldspar, chromian spinel, rutile, andphlogopite, with their grain sizes attaining ~ 200 µm(Supplementary Fig. S2h–j, Supplementary Table S8).
The grain boundaries are locally high in K, which is
attributed to the presence of K-feldspar. None of the olivine nephelinite contains the calcite, rutile, nor phlogopite (cf., Wood 1978). The clinopyroxene occurring in thegrain boundaries sealed with minerals coexists with calcite and has chemical composition distinct from those of
the host olivine nephelinite in terms of CaO content (20wt% vs. 23 wt%), ­Na2O content (0.4–1.1 wt% vs. 0.5 wt%),and ­TiO2 content (1.4–1.9 wt% vs. 2.2 wt%) (Fig. 9c; Supplementary Table S6). These corroborate that the olivine

Page 21 of 28

nephelinite magma is irrelevant to the formation of thegrain boundaries sealed with minerals, and hence weinfer that the crystallization of the grain boundary-filling agent took place before the infiltration of the olivine
nephelinite magma. This along with the presence of calcite in the grain boundaries sealed with minerals suggest
involvement of a carbonaceous fluid before the infiltration of olivine nephelinite magma. The carbonaceous
fluid had probably infiltrated into the xenolith from theouter margin of xenolith, considering that the distribution of grain boundaries sealed with minerals is controlled by the orientation of the outer margin of xenolith(Fig. 2d–i).
The chromian spinel grains in the FMA are crosscut
by the calcite veins (Fig. 3d, j), and the contacting chromian spinel wall contains no ferric iron (Figs. 3k, 9d; Supplementary Table S7). These petrographic and chemical
characteristics also suggest that the AI70 lherzolite wasonce infiltrated by the carbonaceous fluid. The calcite
veins in the chromian spinel orienting at high angles
against elongate spinel morphology (Fig. 3c, d, j) indicate that the carbonaceous fluid was introduced along
brittle fractures developed in the chromian spinel. Since
the garnet has a higher strength than the spinel (Nicolas
and Poirier 1976), the relationship that the only chromianspinel inclusions have the calcite vein cannot be simply
explained. The termination of calcite veins at the boundary between the chromian spinel and the host FMA corroborates that the brittle failure developed during thedecomposition of garnet. Because the volume change by
the garnet-breakdown reaction (2) calculated from molar
volume data listed in Robie (1966) is positive, the competent chromian spinel grains embedded in the garnetmight have been fractured and sucked the carbonaceous
fluid, which precipitated calcite therein (see the panel
5 in Fig. 11b). The area of high concentration of calcite
veins in chromian spinel inclusions is on the extension ofvein (see the white dotted circle in Fig. 3c), from which
Type 2 FMA was developed. This demonstrates that theformation of calcite veins in chromian spinel and Type
2 FMA are penecontemporaneous (see the panels 5 and
6 in Fig. 11b). The decomposition of calcite veins cutting the chromian spinel in the FMA within the 1 cmwide zone from the contact with the olivine nephelinite(see the yellow arrowheads facing each other in Fig. 3c)substantiates that the formation of calcite veins and thedecomposition of garnet into FMA predated the infiltration of olivine nephelinite magma into the xenolith (seethe panel 7 in Fig. 11b).
We speculated above that the isochemical garnet
breakdown reaction (2) occurred during rapid decompression, which might have accelerated the infiltration ofcarbonaceous fluid (see the panel 5 in Fig. 11b). All types

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

of FMA have olivine + plagioclase + spinel assemblage,
but not orthopyroxene + plagioclase + spinel, which are
expected from the isochemical garnet breakdown reaction (2) (Obata et al. 2013). This implies that orthopyroxene + plagioclase + spinel assemblage was replaced byolivine + plagioclase + spinel assemblage during the xenolith ascent.
The Type 2 FMA is higher in Na and LREE than Type
1 FMA (Figs. 2d, 9f; Supplementary Fig. S5). In contrast,
we demonstrated that the bulk chemical compositionof FMA can be explained by decomposition of pyroperich garnet only if alkalis are ignored (Supplementary
Table S11). Therefore, it is inferred that carbonaceous
fluid influx could have triggered the transformation oforthopyroxene + plagioclase + spinel into olivine + plagioclase + spinel due to the introduction of Na enhancing nucleation and growth of sodic plagioclase (Fig. 4) atlower pressures, where olivine and plagioclase are stable(see the black diamond 5 in Fig. 11a).

Page 22 of 28

It is well known that the complexes comprising carbonatite and alkaline silicates show bimodal distributions of
the rock types (e.g., Le Bas 1987). Liquid immiscibility isa favored petrogenetic process to enhance evolved alkaline silicate magma generation in association with carbonatites (e.g., Schmidt and Weidendorfer 2018). Melting
experiments showed that an alkali-rich system tends todrive the carbonate–silicate liquid immiscibility (Lee andWillie 1997; Brooker and Kjarsgaard 2011). Considering that our olivine nephelinite is high in alkali contents(~ 5 wt% ­Na2O + ­K2O; Fig. 6d; Supplementary Table S2),liquid immiscibility could have involved in its origin.
We hypothesize that carbonaceous fluid and olivine
nephelinite magma separated from an alkaline silicate
magma sequentially entrapped the xenolith and transported from the garnet-stability field (Fig. 12) at a highascent rate (0.1 m ­s−1, see Sect. 8.3 for the calculation).
The carbonaceous fluid responsible for the formation of Type 2 FMA may have been exsolved from theascending alkaline silicate magma expelling olivine

Fig. 12 A schematic illustrations of transportation processes of mantle xenoliths from Aitutaki Island. a Alkaline silicate magma was first formedin the asthenospheric mantle and upwelled in response to small-scale sublithospheric convection. Carbonaceous fluid and olivine nephelinite
magma formed from the alkaline silicate magma via liquid immiscibility. b Because of low density of the carbonaceous fluid, it was segregatedin the leading tip of the ascending fluid, leaving olivine nephelinite magma behind. The carbonaceous fluid broke through the mantle
and ascended entrapping mantle xenoliths in the garnet-stability field. Breakdown reaction of pyrope-rich garnet to Type 1 FMA was driven
by the upwelling ascent. c During the carbonaceous fluid transported the xenoliths in a short time, Type 2 FMA was formed by the infiltrationof the carbonaceous fluid. The carbonaceous fluid was not able to hold the xenoliths for a long time because of its low viscosity and low density,
facilitating the descent of the xenoliths toward the boundary with the olivine nephelinite magma. d The xenoliths dropped into the olivine
nephelinite magma, which infiltrated into the xenoliths to form Type 3 FMA. The olivine nephelinite magma eventually brought the xenoliths
to the surface. See Sect. 8.5 for further details on the transportation of AI70 lherzolite xenolith. The black diamonds 4 to 7 correspond to thosein Fig. 11

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

nephelinite component (Fig. 12a). The carbonaceous
fluid may have been quickly segregated and ascentleaving olivine nephelinite magma behind becauseof the low density and viscosity of carbonaceous fluid(~ 0.01 Pa ­s−1; e.g. Kono et al. 2014) and entrapping
mantle xenoliths (Fig. 12b). The carbonaceous fluid
quickly infiltrated into the xenolith owing to its low
viscosity (Fig. 12c), which later frozen along open grainboundaries sealed with minerals. Because of the low
density and viscosity of the carbonaceous fluid, thexenoliths quickly sank through the fluid into the olivine
nephelinite magma ascending behind (Fig. 12c, d). The
olivine nephelinite magma transported the xenoliths
to the surface, during which the olivine nephelinite
magma infiltrated into the xenoliths to form Narich glass-sealing grain boundaries and Type 3 FMA(Fig. 12d).
8.6 Origin of native iron and nepheline in FMA

Native iron and nepheline in FMA are closely associated to each other with direct contact forming composite
grains (Fig. 3f ), indicating that they are genetically related
and formed simultaneously as composite grain. Suchnative iron-nepheline composite grains in Types 1 and 2
FMAs are similar in grain size (~ 5 µm in diameter) andmuch larger than the other constituent minerals (Fig. 3f,h). The native iron-nepheline composite grains are locallyabsent along the FMA grain margins and around the network-like fractures forming dendritic network of Type
2 FMA. These microstructural relationships substantiate that there is no genetic link between the native ironnepheline composite grains and Types 2 and 3 FMAs,
and that they had already been present when Types 2 and3 FMAs formed.
Metallic iron is one of the minerals that can be stable
under reduced conditions below the iron–wüstite (IW)buffer. Yet, such low oxidation state is not expected in theoceanic upper mantle settings (e.g., Eguchi and Dasgupta2018). Experimental studies performed for fertile mantle compositions as analogous to fertile AI70 lherzolite
demonstrated that the metallic iron reaches saturation athigh temperatures (1450–1600 °C) and pressures higher
than 7 GPa (Rohrbach et al. 2007, 2011). In such cases,
redox equilibrium with metallic iron is expected by disproportionation of ­Fe2+ into both ­Fe0 and ­Fe3+ with preferential incorporation of ­Fe3+ into silicate phases, such as
majorite garnet. This phenomenon can account for theorigin of native iron-nepheline composite grains at highpressure, but there has been no report of majoritic garnetwith pervasive distribution of such composite inclusions.
The homogeneous distribution of native iron-nepheline
composite grains in Type 1 FMA suggests a decomposition origin from Fe and Na-rich components in the garnet

Page 23 of 28

under a reducing condition (see the panel 4 in Fig. 11b).
Our native iron is clearly distinct in petrography from
reported native iron along cracks in garnet inclusion
hosted by diamond aggregate (e.g., Jacob et al. 2004). The
native iron and nepheline were later decomposed during
the garnet-breakdown reaction (2) with the introduction of carbonaceous fluid along fractures, which drove
the transformation of Type 1 FMA to Type 2 FMA (seethe panels 5 and 6 in Fig. 11b). The native iron-nepheline
composite grains are absent or depleted around the dendritic veins (Fig. 3i), around the chromian spinel inclusions (Fig. 3j), and in the outer margin of FMA (Fig. 3d),where Type 2 FMA occupies (Fig. 2g). The high Fe content and low Mg content of Type 2 FMA around the dendritic veins, chromian spinel grains, and grain margin of
FMA (Figs. 2d–i) indicate that the native iron-nepheline
composite grains were decomposed therein. The absenceof ­Fe3+ along the chromian spinel rims and calcite veins(Fig. 3j, k) might reflect that the chromian spinel with ~ 5
wt% ­Fe2O3 (Supplementary Table S7) supplied oxygen to
the native iron for oxidation reaction. Therefore, oxygen
fugacity of the carbonaceous fluid is required to havebeen slightly higher than the IW buffer.
8.7 Implications for small‑scale sublithospheric convection

In this subsection, we attempt to decipher the mantle
dynamics before the exhumation of AI70 lherzolite. In
Aitutaki Island, melt depletion ages (i.e., Re-depletion
model age) calculated from 187Os/188Os of the peridotite xenoliths show a weak positive correlation withthe whole-rock ­Al2O3 content except for two outliers(Fig. 6c). The PUM-like fertile xenoliths with high ­Al2O3
content give T­ RD ~ 0 Ga (Fig. 6c). This implies that the fertile lherzolite including the AI70 lherzolite experienced
neither melting beneath the mid-ocean ridge nor significant melting in ancient times anywhere else. The AI70
lherzolite thus does not represent an ancient, depleted
mantle domain, but a fertile mantle close to PUM interms of the bulk major-element and HSE compositions(Fig. 6a, b). It is most plausible that the AI70 lherzolite
represent asthenospheric mantle that underwent negligible melting if any.
Aitutaki Island records old volcanic stage (~ 9.4 Ma)
and young volcanic stage (~ 1.4 Ma) (e.g., Rose and Koppers 2019). Xenoliths are included in lavas of the young
volcanic stage (Fodor et al. 1982). We advocate that amantle upwelling could be a heat source of the youngstage volcanism. The weak heating inferred from the diffusion modeling of large orthopyroxene grain (Fig. 10)
occurred just before the extraction of xenoliths. The weakheating is attributed to segregation and transportation ofparental alkaline silicate magma in the upwelling mantle,

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 24 of 28

Fig. 13 A schematic mantle dynamics model inferred from AI70 lherzolite. Modified after Ballmer et al. (2007). See Sect. 8.7 for further detailsof the inferred small-scale sublithospheric convection roll aligned along the Cook Islands. The black diamonds 4 to 7 correspond to those in Fig. 11.
Not to scalefrom which the olivine nephelinite magma of the young
volcanic stage was derived (Fig. 13).
Since the decompression rate of the mantle estimatedfrom the AI70 lherzolite (~ 0.6 cm ­year−1) is ~ 13 times
slower than the velocity of the Pacific Plate (~ 8 cm ­year−1,
Zahirovic et al. 2015), we envisage that the AI70 lherzolite was sampled from part of slowly convecting mantle
just beneath the lithosphere and asthenosphere boundary when the upwelling mantle impinged the base oflithosphere. The depth of lithosphere and asthenosphereboundary estimated for the Southwest Pacific by seismic
observation (Herath et al. 2022) is similar to the depth
at which the AI70 lherzolite resided shortly before thequick extraction (see the solid brown horizontal line inpressure–temperature space in Fig. 11a). The adiabatically upwelled asthenospheric mantle induced formationof an alkaline silicate magma parental to the host olivine
nephelinite (Fig. 11b).
Garnet including dendritic chromian spinel inclusions
before the decomposition of garnet into FMA is mostplausibly explained by growth of garnet by consumptionof spinel and pyroxenes, which is represented by the reaction (1) proceeding to the right by pressure increase (seethe panels 1 and 2 in Fig. 11b). We here speculate that theasthenospheric mantle is circulating beneath the oceanic
lithosphere. There are two circulation patterns of mantle
flow in the oceanic region: (1) spherical vortex caused bymantle plume (e.g., Campbell and Griffith 1990; Agrusta
et al. 2013), or (2) small-scale sublithospheric convection(SSC) (e.g., Richter and Parsons 1975). As a hotter mantle
plume consists of a large bulbous head and narrow feeder
conduit, a colder thin boundary layer of adjacent mantle
is heated by conduction until its temperature and density
get comparable to the plume source, and then it is forced
to stirred by a recirculating flow within the plume head(Campbell et al. 1989). Along the Cook-Austral volcanic
chain, two age-progressive Macdonald trend and Rurutu
trend with EM (enriched mantle) to HIMU (high µ or
high 238U/204Pb) affinities are identified (Fig. 1a) (Jackson
et al. 2020). Thus, the plume-induced radial flow could
have contributed to the convective motion inferred fromthe AI70 lherzolite. However, if this is the case, the AI70
lherzolite was intensely heated and melted, as the potential temperature of the plume is estimated to be hotter
than the adjacent mantle in the range of 1500–1700 °C[cf., White (2010) and references therein]. However, noneAkizawa et al. Progress in Earth and Planetary Science(2024) 11:38

of intense heating nor concomitant sever melt depletion
are recognized in the AI70 lherzolite.
In the Earth’s upper mantle, SSC can develop as instabilities of the thickened thermal boundary layer grownbeneath old oceanic lithosphere. The convective rolls
are aligned by plate motion (Richter and Parsons 1975).
In contrast to the plume-induced radial flow model discussed above, the SSC model does not require any abnormal heating event. Short-lived (~ 20 Ma) volcanism in
the Cook-Austral volcanic chain is also explained by SSC
model (Ballmer et al. 2007). In conclusion, we argue thatthe fertile asthenospheric mantle in the oceanic region isconvecting in response to SSC (Fig. 13) and generated analkaline silicate magma parental to the olivine nephelinite(see Sect. 8.5). Onset of SSC is classically deduced fromthe timing of flattening (~ 70 Ma) on the oceanic depth
and heat flow profiles against the seafloor age (e.g., Steinand Stein 1992). In this regard, SSC can occur beneath
Aitutaki Island located in mature oceanic lithosphere(~ 107 Ma, Fig. 1b).

9 Conclusions
We examined a lherzolite xenolith included in olivine
nephelinite lava from Aitutaki Island, Cook Islands. The
lherzolite xenolith contains reddish fine-grained mineral aggregates (FMAs), which were classified into three
types based on petrography, microstructure, and chemical characteristics. The FMA grains are surrounded byolivine-rich and pyroxene-poor zone and has clustersof anhedral chromian spinel grains as inclusions. No
chromian spinel occurs outside the FMA grains. The
chromian spinel grains are isolated from each other in
two dimensions, but have clusters with similar crystallographic orientations, suggesting they originallyformed large equant single grains, which were later partly
decomposed to those in irregular three-dimensional
morphology. Bulk major-element and trace-element
compositions of FMA are close to those of pyrope-richgarnet in the garnet-bearing peridotite. These petrographic and geochemical observations corroborate that
the FMA is decomposed pyrope-rich garnet, which wastransformed from an aggregate of aluminous spinel andpyroxenes leaving the olivine margin and chromian spinel inclusions.
Large orthopyroxene and clinopyroxene grains show
slight but clear chemical heterogeneity: increases in
Al, Ca, and Cr from the grain center to the rim in theorthopyroxene grain, and increases in Al and Cr fromthe grain center to the rim with a monotonous increasein Ca in the clinopyroxene grain. The Al and Ca zoning
patterns of orthopyroxene were reproduced by a diffusion model with keeping surface concentration in equilibrium with the other coexisting mineral phases under

Page 25 of 28

decompression with weak heating. The decompression
is estimated to have occurred on a time scale of a few
tenths of million years. Because the reaction of pyroxenes + aluminous spinel → olivine + pyrope-rich garnet
resulted from an increase in pressure, we hypothesizethat the asthenospheric mantle was initially descended
and then ascended by small-scale sublithospheric convection controlled by the oceanic plate motion. The lastpressure–temperature condition shortly before xenolith exhumation was estimated to have been 1140 °C
and 2.12 GPa. This condition is within the garnetstability field, being consistent with the FMA origin
as garnet. It is also consistent with the depth of lithosphere and asthenosphere boundary on the Pacific Plate
determined from seismic observations. The FMA wasformed later during quick extraction by an alkaline silicate magma, which underwent exsolution of carbonaceous fluid leaving the olivine nephelinite magma. The
lherzolite xenolith was first entrapped by the carbonaceous fluid for short period of time and then entrapped
and transported by the olivine nephelinite magma to
the surface at ~ 1.4 Ma.
Abbreviations
FMA	Fine-grained mineral aggregate
HREE	Heavy rare-earth element
HSE	Highly siderophile element
LLSVP	Large low shear velocity province
LREE	Light rare-earth element
MREE	Middle rare-earth element
P–T	Pressure–temperature
PUM	Primitive upper mantle
REE	Rare-earth element
SSC	Small-scale sublithospheric convection
TAS	Total alkali versus silica

Supplementary Information
The online version contains supplementary material available at https://​doi.​
org/​10.​1186/​s40645-​024-​00643-w.
Additional file 1.
Additional file 2.
Acknowledgements
We are grateful to N Ogawa, T Kawakami, and Y Sato for helpful assistance for
EPMA; to H Yoshida for helpful assistance during FE-SEM-EDS analyses. We
appreciate helpful assistance and instruction by S Okazaki, G McCormack,
and J Webb during our stay in Cook Islands. We thank N Puna, V Tangatataia,
B Bishop, T Arere, V Puapii, B Maxwell, A Simona, M Herman, M Rima, and theother staffs in The National Environment Service Cook Islands for assistance of
our field work in Cook islands; and Government of the Cook Islands, Office of
the Prime Minister, Research Permitting Authority, T&M Heather, Charlie Strickland, and Te Mato Vai for allowing us to collect rock samples. We acknowledge
discussions with T Ishii, H Sumino, M Hamada, and I McIntosh. This manuscript
was improved by comments from J Yamamoto and an anonymous reviewer.
Author contributions
NA, TKo, and AI designed the study. NA, AI, CO, and RF had field expeditions at
Aitutaki Island. NA, CO, and AY conducted mineral major-element analyses. NA
and KO did diffusion modeling and numerical simulations using the results ofmineral major-element analyses, and conducted qualitative element analyses

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Page 26 of 28

for FMA. NA, TKo, and AY created elemental-distribution maps using either
EPMA or X-ray analytical microscope. NA, AI, CO, and RF measured wholerock compositions. AM and YI investigated FMA using FE-TEM-EDS. SW andTN carried out EBSD analyses to determine the crystallographic orientationof the chromian spinel grains. Raman spectroscopy was conducted by TKa.
Mineral trace-element compositions were analyzed by AT, TM, and SA. NA,
KO, TKo, and AI wrote basic framework of the manuscript and all authors
contributed to improve the manuscript. All authors read and approved thefinal manuscript.

hotspot found in the Austral Islands. Geology 30:1023–1026. https://​doi.​
org/​10.​1130/​0091-​7613(2002)​030%​3c1023:​ASTMHF%​3e2.0.​CO;2
Bonneville A, Dosso L, Hildenbrand A (2006) Temporal evolution and geochemical variability of the South Pacific superplume activity. Earth Planet
Sci Lett 244:251–269. https://​doi.​org/​10.​1016/j.​epsl.​2005.​12.​037
Boyd FR, Nixon PH (1978) Ultramafic nodules from the Kimberley pipes, South
Africa. Geochim Cosmochim Acta 42:1367–1382. https://​doi.​org/​10.​1016/​
0016-​7037(78)​90042-X
Brey GP, Köhler T (1990) Geothermobarometry in four-phase lherzolites II. New
thermobarometers, and practical assessment of existing thermobarometers. J Petrol 31:1378–1378. https://​doi.​org/​10.​1093/​petro​logy/​31.6.​1353
Brooker RA, Kjarsgaard BA (2011) Silicate-carbonate liquid immiscibility andphase relations in the system ­SiO2–Na2O–Al2O3–CaO–CO2 at 0·1–2·5 GPa
with applications to carbonatite genesis. J Petrol 52:1281–1305. https://​doi.​org/​10.​1093/​petro​logy/​egq081
Bultitude RJ, Green DH (1971) Experimental study of crystal-liquid relationships at high pressures in olivine nephelinite and basanite compositions.
J Petrol 12:121–147. https://​doi.​org/​10.​1093/​petro​logy/​12.1.​121
Campbell IH, Griffiths RW (1990) Implications of mantle plume structure forthe evolution of flood basalts. Earth Planet Sci Lett 99:79–93. https://​doi.​
org/​10.​1016/​0012-​821X(90)​90072-6
Campbell IH, Griffiths RW, Hill RI (1989) Melting in an Archaean mantle plume:
heads it’s basalts, tails it’s komatiites. Nature 339:697–699. https://​doi.​org/​
10.​1038/​33969​7a0
Carswell DA, Dawson JB (1970) Garnet peridotite xenoliths in south Africankimberlite pipes and their petrogenesis. Contrib Miner Petrol 25:163–184.
https://​doi.​org/​10.​1007/​BF003​71129
Cherniak DJ, Dimanov A (2010) Diffusion in pyroxene, mica and amphibole.
Rev Miner Geochem 72:641–690. https://​doi.​org/​10.​2138/​rmg.​2010.​72.​14
Cherniak DJ, Liang Y (2022) Calcium diffusion in enstatite, with application to
closure temperature of the Ca-in-opx thermometer. Geochim Cosmochim Acta 332:124–137. https://​doi.​org/​10.​1016/j.​gca.​2022.​06.​018
Dasgupta R, Hirschmann M (2006) Melting in the Earth’s deep upper mantle
caused by carbon dioxide. Nature 440:659–662. https://​doi.​org/​10.​1038/​
natur​e04612
Day JMD, Walker RJ, Warren JM (2017) 186Os–187Os and highly siderophileelement abundance systematics of the mantle revealed by abyssal
peridotites and Os-rich alloys. Geochim Cosmochim Acta 200:232–254.
https://​doi.​org/​10.​1016/j.​gca.​2016.​12.​013
Demouchy S, Ishikawa A, Tommasi A, Alard O, Keshav S (2015) Characterization of hydration in the mantle lithosphere: Peridotite xenoliths from theOntong Java Plateau as an example. Lithos 212–215:189–201. https://​doi.​
org/​10.​1016/j.​lithos.​2014.​11.​005
Duncan RA, McDougall I (1976) Linear volcanism in French Polynesia. J Volcanol Geotherm Res 1:197–227. https://​doi.​org/​10.​1016/​0377-​0273(76)​90008-1
Eguchi J, Dasgupta R (2018) Redox state of the convective mantle from­CO2-trace element systematics of oceanic basalts. Geochem Perspect
Lett 8:17–21. https://​doi.​org/​10.​7185/​geoch​emlet.​1823
Ezad IS, Dobson DP, Thomson AR, Jennings ES, Hunt SA, Brodholt JP (2022)
Kelyphite textures experimentally reproduced through garnet breakdown in the presence of a melt phase. J Petrol 63:egac110. https://​doi.​
org/​10.​1093/​petro​logy/​egac1​10
Falloon TJ, Green DH (1989) The solidus of carbonated, fertile peridotite.
Earth Planet Sci Lett 94:364–370. https://​doi.​org/​10.​1016/​0012-​821X(89)​90153-2
Fischer KM, Ford HA, Abt DL, Rychert CA (2010) The lithosphere-asthenosphere boundary. Annu Rev Earth Planet Sci 38:551–575. https://​doi.​org/​
10.​1146/​annur​ev-​earth-​040809-​152438
Fodor RV, Bauer GR, Keil K (1982) Ultramafic inclusions and megacrysts inolivine nephelinite, Aitutaki Island, Cook Islands. NZ J Geol Geophys
25:67–76. https://​doi.​org/​10.​1080/​00288​306.​1982.​10422​505
Gasparik T (1987) Orthopyroxene thermobarometry in simple and complex
systems. Contr Miner Petrol 96:357–370. https://​doi.​org/​10.​1007/​BF003​
71254
Gasparik T (2000) An internally consistent thermodynamic model for thesystem CaO–MgO–Al2O3–SiO2 derived primarily from phase equilibrium
data. J Geol 108:103–119. https://​doi.​org/​10.​1086/​314389
Gasparik T (2003) Phase diagrams for geoscientists. Springer, Berlin

Funding
This work was supported by MEXT/JSPS KAKENHI (Japan) Grant Numbers
24K00733, 24K07189, 23H01267, 23K03544, 23H01269, 20KK0079, andJP15H05831.
Availability of data and materials
The datasets supporting the conclusions of this article are all included within
the article and supplementary files.

Declarations
Competing interests
The authors declare that they have no competing interests.
Author details
1
Atmosphere and Ocean Research Institute, University of Tokyo, 5‑1‑5Kashiwanoha, Kashiwa, Chiba 277‑8564, Japan. 2 Department of Environmental Science, Tokyo Gakugei University, 4‑1‑1 Nukuikita, Koganei, Tokyo184‑8501, Japan. 3 Graduate School of Human and Environmental Studies,
Kyoto University, Yoshida‑Nihonmatsu, Sakyo, Kyoto 606‑8501, Japan. 4 Department of Earth and Planetary Sciences, Tokyo Institute and Technology, 2‑12‑1
Ookayama, Meguro, Tokyo 152‑8551, Japan. 5 Division of Earth and Planetary
Sciences, Kyoto University, Kitashirakawa‑Oiwake, Sakyo, Kyoto 606‑8502,
Japan. 6 Department of Earth and Planetary Science, University of Tokyo,
Tokyo 113‑0013, Japan. 7 Department of Geosciences, Faculty of Science,
Shizuoka University, 836 Ohya, Suruga, Shizuoka 422‑8529, Japan. 8 Facultyof Geosciences and Civil Engineering, Kanazawa University, Kakuma, Kanazawa, Ishikawa 920‑1192, Japan.
Received: 13 February 2024 Accepted: 11 June 2024

References
Adam C, Bonneville A (2005) Extent of the South Pacific Superswell. J Geophys
Res 110:B09408. https://​doi.​org/​10.​1029/​2004J​B00346
Agrusta R, Arcay D, Tommasi A, Davaille A, Ribe N, Gerya T (2013) Small-scaleconvection in a plume-fed low-viscosity layer beneath a moving plate.
Geophys J Int 194:591–610. https://​doi.​org/​10.​1093/​gji/​ggt128
Aoki T, Ozawa K, Bodinier JL, Boudier F, Sato Y (2020) Thermal and decompression history of the Lanzo Massif, northern Italy: Implications for thethermal structure near the lithosphere–asthenosphere boundary. Lithos
372–373:105661. https://​doi.​org/​10.​1016/j.​lithos.​2020.​105661
Arai S (1987) An estimation of the least depleted spinel peridotite on the basisof olivine-spinel mantle array, Neues Jahrb. Min Monatsh 1987:347–354
Ballmer MD, van Hunen J, Ito G, Tackley PJ, Bianco TA (2007) Non-hotspot
volcano chains originating from small-scale sublithospheric convection.
Geophys Res Lett 34:L23310. https://​doi.​org/​10.​1029/​2007G​L0316​36
Becker H, Horan MF, Walker RJ, Gao S, Lorand JP, Rudnick RL (2006) Highly
siderophile element composition of the Earth’s primitive upper mantle:
constraints from new data on peridotite massifs and xenoliths. Geochim
Cosmochim Acta 70:4528–4550. https://​doi.​org/​10.​1016/j.​gca.​2006.​06.​004
Bonechi B, Stagno V, Kono Y, Hrubiak R, Ziberna L, Andreozzi GB, Perinelli C,
Gaeta M (2022) Experimental measurements of the viscosity and melt
structure of alkali basalts at high pressure and temperature. Sci Rep
12:2599. https://​doi.​org/​10.​1038/​s41598-​022-​06551-7
Bonneville A, Le Suavé R, Audin L, Clouard V, Dosso L, Gillot PY, Janney P,
Jordahl K, Maamaatuaiahutapu K (2002) Arago Seamount: the missing

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Goto A, Yokoyama K (1988) Lherzolite inclusions in olivine nephelinite tuff
from Salt Lake Crater. Hawaii Lithos 21:67–80. https://​doi.​org/​10.​1016/​
0024-​4937(88)​90006-0
Griffin WL, O’Reilly SY, Ryan CG (1999a) The composition and origin of subcontinental lithospheric mantle. In: Fei Y, Bertka CM, Mysen BO (eds)
Mantle petrology: field observations and high pressure experimentation.
A tribute to Francis R. Geochemical Society, Houston, pp 13–45
Griffin WL, Fisher NI, Friedman J, Ryan GG, O’Reilly SY (1999b) Cr-pyrope
garnets in the lithospheric mantle 1. Compositional systematics andrelations to tectonic setting. J Petrol 40:679–704. https://​doi.​org/​10.​1093/​
petroj/​40.5.​679
Hanyu T, Dosso L, Ishizuka O, Tani K, Hanan BB, Adam C, Nakai S, Senda R,
Chang Q, Tatsumi Y (2013) Geochemical diversity in submarine HIMU
basalts from Austral Islands, French Polynesia. Contrib Miner Petrol
166:1285–1304. https://​doi.​org/​10.​1007/​s00410-​013-​0926-x
Harigane Y, Mizukami T, Morishita T, Michibayashi K, Abe N, Hirano N (2011)
Direct evidence for upper mantle structure in the NW Pacific Plate: microstructural analysis of a petit-spot peridotite xenolith. Earth Planet Sci Lett
302:194–202. https://​doi.​org/​10.​1016/j.​epsl.​2010.​12.​011
Herath P, Stern TA, Savage MK, Bassett D, Henrys S (2022) Wide-angle seismicreflections reveal a lithosphere-asthenosphere boundary zone in the
subducting Pacific Plate, New Zealand. Sci Adv 8:eabn5697. https://​doi.​
org/​10.​1126/​sciadv.​abn56​97
Herzberg C, Asimow PD (2015) PRIMELT3 MEGA.XLSM software for primarymagma calculation: Peridotite primary magma MgO contents from theliquidus to the solidus. Geochem Geophys Geosyst 16:563–578. https://​doi.​org/​10.​1002/​2014G​C0056​31
Hirschmann MM (2000) Mantle solidus: experimental constraints andthe effect of peridotite composition. Geochem Geophys Geosyst
1:2000GC000070. https://​doi.​org/​10.​1029/​2000G​C0000​70
Ibarguchi JIG, Ábalos B, Azcarraga J, Puelles P (1999) Deformation, highpressure metamorphism and exhumation of ultramafic rocks in a deep
subduction/collision setting (Cabo Ortegal, NW Spain). J Metamorph
Geol 17:747–764. https://​doi.​org/​10.​1046/j.​1525-​1314.​1999.​00227.x
Ishikawa A, Maruyama S, Komiya T (2004) Layered lithospheric mantle beneath
the Ontong Java Plateau: implications from xenoliths in alnöite, Malaita,
Solomon Islands. J Petrol 45:2011–2044. https://​doi.​org/​10.​1093/​petro​logy/​egh046
Ishikawa A, Pearson DG, Dale CW (2011) Ancient Os isotope signatures fromthe Ontong Java Plateau lithosphere: tracing lithospheric accretion
history. Earth Planet Sci Lett 301:159–170. https://​doi.​org/​10.​1016/j.​epsl.​
2010.​10.​034
Jackson MG, Halldórsson SA, Price A, Kurz MD, Konter JG, Koppers AAP, Day
JMD (2020) Contrasting old and young volcanism from Aitutaki, Cook
Islands: implications for the origins of the Cook-Austral volcanic chain. J
Petrol 61:egaa037. https://​doi.​org/​10.​1093/​petro​logy/​egaa0​37
Jacob DE, Kronz A, Viljoen KS (2004) Cohenite, native iron and troilite inclusions in garnets from polycrystalline diamond aggregates. Contrib Miner
Petrol 146:566–576. https://​doi.​org/​10.​1007/​s00410-​003-​0518-2
Kawakatsu H, Kumar P, Takei Y, Shinohara M, Kanazawa T, Araki E, Suyehiro K(2009) Seismic evidence for sharp lithosphere-asthenosphere boundaries
of oceanic plates. Science 324:499–502. https://​doi.​org/​10.​1126/​scien​ce.​11694​99
Kawakatsu H, Utada H (2017) Seismic and electrical signatures of thelithosphere–asthenosphere system of the normal oceanic mantle.
Annu Rev Earth Planet Sci 45:139–167. https://​doi.​org/​10.​1146/​annur​
ev-​earth-​063016-​020319
Klemme S, O’Neill HS (2000) The near-solidus transition from garnet lherzoliteto spinel lherzolite. Contrib Miner Petrol 138:237–248. https://​doi.​org/​10.​
1007/​s0041​00050​560
Kono Y, Kenney-Benson C, Hummer D, Ohfuji H, Park C, Shen G, Wang Y, Kavner
A, Manning CE (2014) Ultralow viscosity of carbonate melts at high pressures. Nat Commun 5:5091. https://​doi.​org/​10.​1038/​ncomm​s6091
Kushiro I, Yoder HS (1966) Anorthite–forsterite and anorthite—enstatite reactions and their bearing on the basalt–eclogite transformation. J Petrol
7:337–362. https://​doi.​org/​10.​1093/​petro​logy/7.​3.​337
Le Bas MJ (1987) Nephelinites and carbonatites. In: Fitton JG, Upton BGJ (eds)
Alkaline igneous rocks, vol 30. The Geological Society of London, Piccadilly, pp 13–45. https://​doi.​org/​10.​1144/​GSL.​SP.​1987.​030.​01.0

Page 27 of 28

Le Bas MJ, Le Maitre RW, Streckeisen A, Zanettin B (1986) A chemical classification of volcanic rocks based on the total alkali–silica diagram. J Petrol
27:745–750. https://​doi.​org/​10.​1093/​petro​logy/​27.3.​745
Le Maitre RW, Streckeisen A, Zanettin B, Le Bas MJ, Bonin B, Bateman P (2002)
Igneous rocks: a classification and glossary of terms: recommendationsof the international union of geological sciences subcommission on thesystematics of igneous rocks, 2nd edn. Cambridge University Press, NewYork
Lee W, Wyllie PJ (1997) Liquid immiscibility between nepherinite and carbonatite from 1.0 to 2.5 GPa compared with mantle melt compositions.
Contrib Miner Petrol 127:1–16. https://​doi.​org/​10.​1007/​s0041​00050​261
Litasov K, Taniguchi H (2002) Mantle evolution beneath the Baikal Rift. CNEAS
monograph series, vol 5. Center for Northeast Asian Studies, Tohoku
University
McDonough WF, Sun SS (1995) The composition of the earth. Chem Geol
120:223–253. https://​doi.​org/​10.​1016/​0009-​2541(94)​00140-4
McNutt MK, Judge AV (1990) The superswell and mantle dynamics beneaththe South Pacific. Science 248:969–975. https://​doi.​org/​10.​1126/​scien​ce.​
248.​4958.​969
Moukadiri A, Kornprobst J (1984) Garnet and/or spinel bearing pyroxenites
in alkaki basalts near Azrou (Middle Atlas, Morocco): mantle derived
alumina-rich xenoliths related to the “ariegite-grospydite” trend. Devel
Petrol 11:179–189. https://​doi.​org/​10.​1016/​B978-0-​444-​42274-3.​50021-4
Müller RD, Zahirovic S, Williams SE, Cannon J, Seton M, Bower DJ, Tetley MG,
Heine C, Breton EL, Liu S, Russell SHJ, Yang T, Leonard J, Gurnis M (2019) A
global plate model including lithospheric deformation along major rifts
and orogens since the Triassic. Tectonics 38:1884–1907. https://​doi.​org/​
10.​1029/​2018T​C0054​62
Neal CR, Nixon PH (1985) Spinel-garnet relationships in mantle xenoliths from
the Malaita Alnöites, Solomon Islands, South-Western Pacific. Trans Geol
Soc S Afri 88:347–354
Nickel KG (1986) Phase equilibria in the system ­SiO2–MgO–Al2O3–CaO–Cr2O3(SMACCR) and their bearing on spinel/garnet lherzolite relationships.
Neues Jahrbuch Miner Abh 155:259–287
Nickel KG, Green DH (1985) Empirical geothermobarometry for garnet
peridotites and implications for the nature of the lithosphere, kimberlites
and diamonds. Earth Planet Sci Lett 73:158–170. https://​doi.​org/​10.​1016/​
0012-​821X(85)​90043-3
Nicolas A, Poirier JP (1976) Crystalline plasticity and solid state flow in metamorphic rocks. Wiley, New YorkNiu Y (2004) Bulk-rock major and trace element compositions of abyssal
peridotites: implications for mantle melting, melt extraction and postmelting processes beneath mid-ocean ridges. J Petrol 45:2423–2458.
https://​doi.​org/​10.​1093/​petro​logy/​egh068
Niu Y, Hékinian R (1997) Basaltic liquids and harzburgitic residues in the Garrett
Transform: a case study at fast-spreading ridges. Earth Planet Sci Lett
146:43–258. https://​doi.​org/​10.​1016/​S0012-​821X(96)​00218-X
Obata M, Ozawa K, Naemura K, Miyake A (2013) Isochemical breakdown ofgarnet in orogenic garnet peridotite and its implication to reaction kinetics. Miner Petrol 107:881–895. https://​doi.​org/​10.​1007/​s00710-​012-​0260-4
Obata M, Ohi S, Miyake A (2014) Experimental synthesis of isochemical
kelyphite—a preliminary report. J Miner Petrol Sci 109:91–96. https://​doi.​
org/​10.​2465/​jmps.​13102​2a
O’Hara MJ, Richardson SW, Wilson G (1971) Garnet-peridotite stability andoccurrence in crust and mantle. Contrib Miner Petrol 32:48–68. https://​doi.​org/​10.​1007/​BF003​72233
Ozawa K (1997) P-T history of an ascending in orthopyroxene: a case Hokkaido,
northern Japan. Mem Geol Soc Jpn 47:107–122
Ozawa K (2004) Thermal history of the Horoman peridotite complex: a recordof thermal perturbation in the lithospheric mantle. J Petrol 45:253–273.
https://​doi.​org/​10.​1093/​petro​logy/​egg110
Paquet M, Day JMD, Brown DB, Walter CL (2022) Effective global mixing of thehighly siderophile elements into Earth’s mantle inferred from oceanic
abyssal peridotites. Geochm Cosmochim Acta 316:347–362. https://​doi.​
org/​10.​1016/j.​gca.​2021.​09.​033
Regelous M, Weinzierl CG, Haase CM (2015) Controls on melting at spreading ridges from correlated abyssal peridotite—mid-ocean ridge basalt
compositions. Earth Planet Sci Lett 449:1–11. https://​doi.​org/​10.​1016/j.​
epsl.​2016.​05.​017

Akizawa et al. Progress in Earth and Planetary Science(2024) 11:38

Richardson SH, Erlank AJ, Hart SR (1985) Kimberlite-borne garnet peridotite
xenoliths from old enriched subcontinental lithosphere. Earth Planet Sci
Lett 75:116–128. https://​doi.​org/​10.​1016/​0012-​821X(85)​90094-9
Richter FM, Parsons B (1975) On the interaction of two scales of convection in
the mantle. J Geophys Res 80:2529–2541. https://​doi.​org/​10.​1029/​JB080​
i017p​02529
Robie RA (1966) Thermodynamic properties of minerals. In: Sydney P, Clark P(eds) Handbook of physical constants, GSA memoir, vol 97. GeologicalSociety of America, New York
Rohrbach A, Ballhaus C, Golla-Schindler U, Ulmer P, Kamenetsky VS, Kuzmin DV(2007) Metal saturation in the upper mantle. Nature 449:456–458. https://​doi.​org/​10.​1038/​natur​e06183
Rohrbach A, Ballhaus C, Ulmer P, Golla-Schindler U, Schönbohm D (2011)
Experimental evidence for a reduced metal-saturated upper mantle. J
Petrol 52:717–731. https://​doi.​org/​10.​1093/​petro​logy/​egq101
Rose J, Koppers AAP (2019) Simplifying age progressions within the CookAustral Islands using ARGUS-VI high-resolution 40Ar/39Ar incremental
heating ages. Geochem Geophys Geosyst 20:4756–4778. https://​doi.​org/​
10.​1029/​2019G​C0083​02
Russell JK, Jones TJ (2023) Transport and eruption of mantle xenoliths creates alagging problem. Commun Earth Environ 4:177. https://​doi.​org/​10.​1038/​
s43247-​023-​00843-0
Ryan WBF, Carbotte SM, Coplan JO, O’Hara S, Melkonian A, Arko R, Weissel RA,
Ferrini V, Goodwillie A, Nitsche F, Bonczkowski J, Zemsky R (2009) Global
multi-resolution topography synthesis. Geochem Geophys Geosyst
10:Q03014. https://​doi.​org/​10.​1029/​2008G​C0023​32
Rychert CA, Harmon N, Constable S, Wang S (2020) The nature of thelithosphere-asthenosphere boundary. J Geophys Res: Solid Earth
125:e2018JB016463. https://​doi.​org/​10.​1029/​2018J​B0164​63
Schmidt MW, Weidendorfer D (2018) Carbonatites in oceanic hotspots. Geology 46:435–438. https://​doi.​org/​10.​1130/​G39621.1
Sen G, Keshav S, Bizimis M (2005) Hawaiian mantle xenoliths and magmas: composition and thermal character of the lithosphere. Am Min
90:871–887. https://​doi.​org/​10.​2138/​am.​2005.​1731
Smith D, Barron BR (1991) Pyroxene-garnet equilibration during cooling in the
mantle. Am Miner 76:1950–1963
Snortum E, Day JMD, Jackson MG (2019) Pacific lithosphere evolution inferredfrom Aitutaki mantle xenoliths. J Petrol 60:1753–1772. https://​doi.​org/​10.​1093/​petro​logy/​egz047
Spera FJ (1984) Carbon dioxide in petrogenesis III: role of volatiles in the ascentof alkaline magma with special reference to xenolith-bearing mafic lavas.
Contr Miner Petrol 88:217–232. https://​doi.​org/​10.​1007/​BF003​80167
Stein CA, Stein S (1992) A model for the global variation in oceanic depth andheat flow with lithospheric age. Nature 359:123–129. https://​doi.​org/​10.​1038/​35912​3a0
Tommashi A, Ishikawa A (2014) Microstructures, composition, and seismic
properties of the Ontong Java Plateau mantle root. Geochem Geophys
Geosyst 15:e2014GC005452. https://​doi.​org/​10.​1002/​2014G​C0054​52
Torsvik TH, Smethurst MA, Burke K, Steinberger B (2006) Large igneous provinces generated from the margins of the large low-velocity provinces inthe deep mantle. Geophys J Int 167:1447–1460. https://​doi.​org/​10.​1111/j.​
1365-​246X.​2006.​03158.x
Tsuchiyama A (1986) Melting and dissolution kinetics: application to partial
melting and dissolution of xenoliths. J Geophys Res 91:9395–9406.
https://​doi.​org/​10.​1029/​JB091​iB09p​09395
Turner DL, Jarrard RD (1982) K-Ar dating of the Cook-Austral island chain: a testof the hot-spot hypothesis. J Volcanol Geotherm Res 12:187–220. https://​doi.​org/​10.​1016/​0377-​0273(82)​90027-0
Yasumoto A, Yoshida K, Kuwatani T, Nakamura D, Svojtka M, Hirajima T (2018)
A rapid and precise quantitative electron probe chemical mapping
technique and its application to an ultrahigh-pressure eclogite from the
Moldanubian Zone of the Bohemian Massif (Nové Dvory, Czech Republic). Am Min 103:1690–1698. https://​doi.​org/​10.​2138/​am-​2018-​6323C​CBY
Yokoyama K (1980) Nikubuchi peridotite body in the sanbagawa metamorphic belt; thermal history of the ‘Al-pyroxene-rich suite’ peridotite body inhigh pressure metamorphic terrain. Contr Miner Petrol 73:1–13. https://​doi.​org/​10.​1007/​BF003​76256
White WM (2010) Oceanic island basalts and mantle plumes: the geochemical
perspective. Annu Rev Earth Planet Sci 38:133–160. https://​doi.​org/​10.​1146/​annur​ev-​earth-​040809-​152450

Page 28 of 28

Wood CP (1978) Petrology of Aitutaki, Cook Islands (Note). N Z J Geol Geophys
21:761–765. https://​doi.​org/​10.​1080/​00288​306.​1978.​10425​205
Zahirovic S, Müller RD, Seton M, Flament N (2015) Tectonic speed limits fromplate kinematic reconstructions. Earth Planet Sci Lett 418:40–52. https://​doi.​org/​10.​1016/j.​epsl.​2015.​02.​037
Ziberna L, Klemme S, Nimis P (2013) Garnet and spinel in fertile and depleted
mantle: insights from thermodynamic modelling. Contrib Miner Petrol
166:411–421. https://​doi.​org/​10.​1007/​s00410-​013-​0882-5

Publisher’s NoteSpringer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.