New evidence of mantle heterogeneity beneath the

Lithos 132-133 (2012) 70–81
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Lithos
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New evidence of mantle heterogeneity beneath the Hyblean Plateau (southeast Sicily,
Italy) as inferred from noble gases and geochemistry of ultramafic xenoliths
A. Correale a,⁎, M. Martelli b, A. Paonita b, A. Rizzo b, L. Brusca b, V. Scribano c
a
b
c
Dipartimento di Scienze della Terra e del Mare (DiSTeM), Università degli Studi di Palermo, Via Archirafi 36, Palermo 90123, Italy
Istituto Nazionale di Geofisica e Vulcanologia, Sezione di Palermo, Via Ugo La Malfa 153, Palermo 90146, Italy
Dipartimento di Scienze Geologiche, Università degli Studi di Catania, Corso Italia 55, Catania 95129, Italy
a r t i c l e
i n f o
Article history:
Received 14 July 2011
Accepted 4 November 2011
Available online 11 November 2011
Keywords:
Xenolith
Mantle
Hyblean Plateau
Metasomatism
Noble gas
Fluid inclusion
a b s t r a c t
We analyzed major and trace elements, Sr and Nd isotopes in ultramafic xenoliths in Miocenic age Hyblean
diatremes, along with noble gases of CO2-rich fluid inclusions hosted in the same products. The xenoliths
consist of peridotites and pyroxenites, which are considered to be derived from the upper mantle. Although
the mineral assemblage of peridotites and their whole-rock abundance of major elements (e.g.,
Al2O3 = 0.8–1.5 wt.%, TiO2 = 0.03–0.08 wt.%) suggest a residual character of the mantle, a moderate enrichment in some incompatible elements (e.g., LaN/YbN = 9–14) highlights the presence of cryptic metasomatic
events. In this context a deep silicate liquid is considered the metasomatizing agent, which is consistent
with the occurrence of pyroxenites as veins in peridotites. Both the Zr/Nb and 143Nd/ 144Nd ratios of the investigated samples reveal two distinct compositional groups: (1) peridotites with Zr/Nb ≈ 4 and 143Nd/
144
Nd ≈ 0.5129, and (2) pyroxenites with Zr/Nb ≈ 20 and 143Nd/144Nd ≈ 0.5130. The results of noble-gas analyses also highlight the difference between the peridotite and pyroxenite domains. Indeed, the 3He/4He and
4
He/ 40Ar* ratios measured in the fluid inclusions of peridotites (respectively 7.0–7.4 ± 0.1 Ra and 0.5–8.2,
where Ra is the atmospheric 3He/4He ratio of 1.38 × 10 − 6) were on average lower than those for the pyroxenites (respectively 7.2–7.6 Ra and 0.62–15). This mantle heterogeneity is interpreted as resulting from a
mixing between two end-members: (1) a peridotitic layer with 3He/4He ≈ 7 Ra and 4He/40Ar* ≈ 0.4, which
is lower than the typical mantle ratio (~ 1–4) probably due to melt extraction events, and (2) metasomatizing
mafic silicate melts that gave rise to pyroxenites characterized by 3He/4He ≈ 7.6 Ra, with a variable 4He/ 40Ar*
due to degassing processes connected with the ascent of magma at different levels in the peridotite wall rock.
The complete geochemical data set also suggests two distinct mantle sources for the xenolithic groups
highlighted above: (1) a HIMU (high-μ)-type source for the peridotites and (2) a DM (depleted mantle)type source for the pyroxenites.
© 2011 Elsevier B.V. All rights reserved.
1. Introduction
Mantle xenoliths from diatremes often exhibit a pristine character
due to the relatively low eruptive temperature and high ascent velocity of such volatile-rich volcanic systems. Thus, a careful investigation
of these xenoliths can provide unique information on upper-mantle
composition and the processes that may modify it (e.g., Beccaluva et
al., 2004; Downes, 2007; Dunai and Baur, 1995; Gautheron et al.,
2005; Vaselli et al., 1995; Zangana et al., 1999). The Central Mediterranean area provides an attractive example of such an approach in a
geodynamically complex region, where the characteristics of the lithospheric mantle have mostly been inferred from geophysical data
(e.g., Berry and Knopoff, 1967; Calcagnile et al., 1982; Finetti and
⁎ Corresponding author at: Dipartimento di Scienze della Terra e del Mare (DiSTeM),
Università degli Studi di Palermo, Via Archirafi 36, Palermo 90123, Italy. Tel.: +39 91
6809273; fax: +39 91 6809449.
E-mail address: [email protected] (A. Correale).
0024-4937/$ – see front matter © 2011 Elsevier B.V. All rights reserved.
doi:10.1016/j.lithos.2011.11.007
Morelli, 1973). Our work within this area focuses on investigating
the mantle beneath the Hyblean Plateau (southeastern Sicily, Italy),
which is one of the rare European volcanic regions where xenoliths
occur.
The Hyblean Plateau has been characterized by several distinct episodes of magmatism, starting from Triassic and lasting until Quaternary (Carbone and Lentini, 1981; Cristofolini, 1966; Rocchi et al.,
1998). Some of the volcanic events brought to the surface a large
number of mantle-derived xenoliths, mainly spinel-facies peridotites
and subordinate pyroxenites. These products have been widely investigated by many authors (Bianchini et al., 2010; Perinelli et al., 2008;
Sapienza and Scribano, 2000; Sapienza et al., 2005; Tonarini et al.,
1996), who highlighted the occurrence of metasomatic events affecting the local mantle. Sr- and Nd-isotope data of Hyblean peridotites
( 87Sr/ 86Sr = 0.70288–0.70309 and 143Nd/ 144Nd = 0.51287–0.51292)
reveal a HIMU (high-μ)-like affinity, while the data for pyroxenites
( 87Sr/ 86Sr = 0.70305–0.70326 and 143Nd/ 144Nd = 0.51292–0.51299),
which differ slightly from those of peridotites, has isotope
A. Correale et al. / Lithos 132-133 (2012) 70–81
characteristics overlapping the alkaline lavas, implying that the pyroxenite domain contributed to the genesis of the Hyblean magmas
(Bianchini et al., 2010; Tonarini et al., 1996).
Geochemical investigation of CO2-rich fluid inclusions hosted in
olivines and pyroxenes confirmed a mantle-derived origin for the ultramafic xenoliths (Sapienza et al., 2005). Microthermometric analyses show entrapment pressures of fluid inclusions in the range
0.75–0.95 GPa, corresponding to a depth of 27–35 km, where spinel
peridotites would be stable (Sapienza et al., 2005). Also, He-isotope
measurements of the fluid inclusions hosted in peridotite minerals
(both olivine and pyroxenes) have values of 7.3 ± 0.3 Ra (where Ra
is the 3He/ 4He ratio of 1.38 × 10 − 6 as measured in air) (Sapienza et
al., 2005), similar to the isotope signature of Pantelleria Island
(Martelli et al., 2008; Parello et al., 2000), and are consistent with a
depleted-mantle (DM) signature of the local mantle. Such values are
the highest of Plio-Quaternary Italian magmatism (e.g., Martelli et
al., 2008) and underline the importance of the Hyblean province in
the evolution of the Italian area.
This study performed a comprehensive investigation of the geochemistry of ultramafic xenoliths from the Hyblean area. Sapienza
et al. (2005) investigated helium in fluid inclusions of Hyblean peridotites, whereas in the present study we investigated both helium
and argon and not only on the peridotites samples but also on a
suite of pyroxenites. The same samples have been analyzed for trace
elements (in whole rocks and clinopyroxenes) and Sr and Nd isotopes. These data give new insights into the different roles played
by peridotites and pyroxenites in determining the Hyblean mantle
characteristics, as well on the contributions of metasomatic
processes.
2. Geological setting
The investigated area is in a critical geodynamical setting characterized by the collision between the European and African plates
(Fig. 1; Barberi et al., 1974), with the Hyblean Plateau located in the
undeformed northern portion of the Pelagian Block, in the foreland
area (Lentini et al., 1996). However, there is a scientific debate
about the nature of the lithosphere beneath the Hyblean region. Indeed, the hypothesis supported by Vai (1994) about the possible oceanic character of this crust contrasts with the more common
geological models that consider this lithospheric block to be in
Fig. 1. Map of Hyblean Plateau. The enlarged area shows the provenance of main xenolithic samples.
71
continuity with the African plate, thus suggesting a continental character (Burollet et al., 1978).
Discontinuous volcanic activity characterized the Hyblean Plateau
from Cretaceous to Pleistocene (Cristofolini, 1966). The products of
the numerous eruptions interrupted the Meso-Cenozoic deep-water
carbonate deposits and the Neogene–Quaternary clastic sequences
(Bianchi et al., 1987). Although Cristofolini (1966) detected a Triassic
igneous layer via drill holes near Ragusa, the oldest eruptive rocks
that outcrop are Cretaceous alkali basalts and are located in the eastern part of the area (Capo Passero, Siracusa, and Augusta; Amore et
al., 1988). After the Cretaceous activity, the volcanism stopped for
about 50 Ma and then restarted during the Miocene age with alkaline
affinity lavas, which can be found in the central-northern area of the
Hyblean Plateau, the so-called volcanic plateau (e.g., Bianchi et al.,
1987). The last eruptive episode, during Plio-Pleistocene, was characterized by the eruption of tholeiite and alkaline basalts and minor
nephelinites (Beccaluva et al., 1998). Some diatreme-related deposits
of Miocene age (Carbone and Lentini, 1981) and some Quaternary
basanitic and nephelinitic lavas sampled a part of lithosphere, carrying a huge amount of ultramafic xenoliths to surface (Scribano,
1987a, 1987b). Among these, the Miocene Valle Guffari diatreme is
characterized by the greatest variety and quantity of deep xenoliths,
and is also the area where most of the investigated samples were collected for the present study (Fig. 1).
3. Analytical techniques
The studied samples consist of peridotitic and pyroxenitic xenoliths found in some Miocene diatremes of the Hyblean area, in particular the Valle Guffari diatreme and Cozzo Molino pipe (Fig. 1). The
samples were selected on the basis of their size (~5 cm for the peridotites and ~ 10 cm for the pyroxenites) and freshness.
Major- and trace-element analyses were performed on both
whole rocks and on selected olivine and pyroxene grains. Whole
rocks were analyzed at the laboratory of SGS Canada using ICP-AES
and ICP-MS technical procedures. Selected portions of samples were
crushed and powdered with an agate mortar, then a weighted aliquot
(~0.10 g) is digested by fusion with sodium peroxide in graphite crucibles or dissolution by multi-acid digestion using a combination of
HCl, HNO3, HF and HClO4. During digestion each sample was split
into two aliquots for ICP-OES and ICP-MS analyses. The accuracy of
the method was determined by analyzing certified reference materials, while its precision was determined with replicate analyses
(and found to be generally better than 10%).
Single pyroxene and olivine crystals were analyzed for major elements using a LEO™ 440 Scanning Electron Microscope coupled to an
Oxford-Link Energy Dispersive Spectroscopy system hosted at the
DiSTeM laboratory, University of Palermo. More details on the procedure of sample preparation and on the analytical technique can be
found in Lopez et al. (2006). Trace-element analyses were performed
using the laser ablation ICP-MS technique at Istituto Nazionale di
Geofisica e Vulcanologia (INGV), Palermo. Selected samples were incorporated into an epoxy-resin puck that was polished before analysis. The analytical system consisted of an Agilent-7500 CX quadrupole
mass spectrometer coupled with an ArF excimer laser ablation system
(GeoLas Pro). During analysis, samples were maintained in a helium
atmosphere, with a laser output energy of 10 J/cm 2, a repetition rate
of 10 Hz, and a 130-μm-diameter circular spot. We used Ca, Si, and
Fe as internal standards and NIST 612 as an external standard. The
NIST 612 analyses were carried out at the start, middle, and end of
each analytical session. The precision was determined during each
analysis session from the variance of ~15 NIST 612 measurements,
which gave a relative standard deviation of b5%. The accuracy, calculated using the BCR-2 international standard, was b10% for most of
the elements.
72
A. Correale et al. / Lithos 132-133 (2012) 70–81
Sr- and Nd-isotope compositions of separated clinopyroxene phenocrysts (typically 1–3 g) were determined at INGV, Osservatorio
Vesuviano Napoli, by thermal ionization MS. Mineral samples that
had been careful hand-picked were crushed to powder in an agate
mortar in order to prepare them for isotope analysis. More detailed
information on the samples preparation and analytical procedures
can be found in Arienzo et al. (2009).
Noble gases were analyzed at the INGV Palermo laboratory by
single-step in-vacuo crushing at a pressure of about 20 MPa (so as
to minimize the contribution of noble gases from the crystal lattice)
coupled with MS. He and Ne were analyzed by GVI-Helix SFT MS,
while Ar was analyzed by a GVI-Argus device. Each sample was analyzed twice, and in each analysis we used about 2 g of olivines and
0.5–1 g of pyroxenes. We followed the same sample preparation
and analytical techniques reported in Nuccio et al. (2008) and
Martelli et al. (2011).
4. Petrography, and bulk-rock and mineral chemistry
On the basis of their mineralogical modal composition (50–75% olivine, 8–25% orthopyroxene, 1–8% clinopyroxene, and 1–3% spinel),
peridotites are classified as anhydrous spinel-facies lherzolite (XIH3 sample) and harzburgite (XIH-1 and XIH-2 samples; Fig. 2). They
have a variable texture, from protogranular to porphyroclastic
(Fig. 3). Data on the major elements are reported in Table 1. Olivines
show an average composition of Fo90, orthopyroxenes were Fs9.4–11.2,
Wo0.9–2.9, and En85.9–89.8, whereas the clinopyroxenes (Cr-diopsides)
comprise Fs4.3–7.2, Wo43.1–49.5, and En46.2–50.1. The results of our analyses are in accordance with available data in the literature for
Hyblean nodules, confirming a general homogeneity in the major
chemistry of the investigated xenoliths (Atzori et al., 1999;
Bianchini et al., 2010; Nimis, 1998; Perinelli et al., 2008; Tonarini et
al., 1996). The peridotite samples are essentially fresh, but sometimes
it was possible to observe a slight degree of serpentinization along the
cracks within olivines or along grain boundaries. Cr–Al spinel was
also present both as interstitial grains and as vermiform intergrowths
with the pyroxene. Kink banding was ubiquitous in the olivine crystals (Fig. 3). Several samples were characterized by local modal increases in pyroxene contents, or centimeter-sized websterite veins.
Four samples could be characterized as pyroxenites based on their
olivine, orthopyroxene, and clinopyroxene percentages (Fig. 2). In
particular, samples XIP-4 and XIP-14 were clinopyroxenites (≥75%
clinopyroxene and ≤ 5% orthopyroxene), while samples XIP-28 and
XIP-17 were websterites (62–70% clinopyroxene and 13–23% orthopyroxene). Clinopyroxene from samples XIP-14, XIP-4, and XIP-17
was an Al-diopside (Al2O3 = 6.8–9.5 wt.%; see Table 1) characterized
by several exsolution lamellae of Ca-poor pyroxene and Al-spinel.
On the other hand, clinopyroxene from XIP-28 and XIC-26 was a Crdiopside (Cr = 0.75 and 1.14 wt.%, respectively; Table 2). The orthopyroxene composition varies in the range Fs9.5–20, Wo0.7–2.9, and
En77.3–89.6.
Sample XIP-28 contains ~ 8 vol.% Fo90 olivine. Most of the pyroxenite samples contain variable amounts of Al–Cr spinel, which is
particularly abundant in sample XIP-4. This explains the exceptionally low silica content (SiO2 = 25.4 wt.%) and high alumina content
(Al2O3 = 23.8 wt.%) in this sample (cf. Table 1). However, it must
be noted that the distribution of the spinel in these xenoliths, and
hence its grain size, were quite irregular. Considering that xenoliths
are fragments of deep rocks, the percentage values therein might not
be representative of the original rock, especially for those with
coarse grain size. In fact, the averaged contents of this particular pyroxenite type deduced previously were 85% clinopyroxene, 5%
orthopyroxene, and 10 vol.% Al-spinel (Punturo and Scribano,
1997). It is also noteworthy that spinel is generally rimmed by a
keliphytized garnet in sample XIP-17 (Table 1). In addition, we considered a first-size composite xenolith (sample XIC-26), consisting
of a harzburgitic peridotite frame cross-cut by two irregular 0.5cm-wide clinopyroxenite veins. The data from the whole-rock analysis reported in Table 1 represent the average composition of this
composite xenolith.
Observations of both peridotites and pyroxenites under the optical
microscope identified array of secondary fluid inclusions in olivine
and pyroxene crystals (Fig. 3). Fluid inclusions were not distributed
uniformly among the different mineralogical phases, in accordance
with previous observations in peridotite paragenesis by Sapienza et
al. (2005). The clinopyroxenes are systematically richer in fluid inclusions relative to coexisting olivines and orthopyroxenes, as also observed in mantle xenoliths from different areas (Porcelli et al., 1986).
Ol
Dunite
XIH-1
XIH-3
Lherzolite
e
Ha
rz b
rlit
e
Wh
urg
ite
XIH-2
Olivine bearing Websterite
Clinopyroxenite
Ortopyroxenite
XIP-28
Opx
Websterite
XIP-17
XIP-4
Fig. 2. Modal composition of the investigated xenoliths.
XIP-14
Cpx
A. Correale et al. / Lithos 132-133 (2012) 70–81
Fig. 3. Thin-section photomicrographs showing typical petrographic features of the peridotite samples: (a) rock-forming minerals (OL, olivine; Opx, orthopyroxene; Cpx,
clinopyroxene; Spl, spinel) and their textural relations (sample XIH-3, crossed polars);
(b) part of a kink-banded olivine grain cross-cut by composite serpentine and carbonate veins (sample XIH-3, crossed polars); and (c) fluid inclusions array within an olivine grain (plane-polarized light).
5. Trace-element and Sr- and Nd-isotope geochemistry
Trace-element data for the peridotite and pyroxenite samples are
listed in Table 2. Fig. 4 shows the chondrite-normalized REE distribution for both whole rocks and clinopyroxenes of peridotites. The REE
patterns are similar in Fig. 4a and b, although the clinopyroxenes
show clear REE enrichments relative to whole rocks, which are due
to the high affinity of REE for the pyroxene structure (Eggins et al.,
1998).
The plots show consistent patterns of both whole rocks and clinopyroxenes among different samples, suggesting a homogeneous
source composition. Compared to chondrite, all samples show evident LREE enrichment (Lan/Ybn ≈ 20) while HREE is slightly depleted
(Fig. 4a). Evidence for this can also be found in previously published
data for other Hyblean peridotites and were attributed to a pervasive
or, more likely, cryptic metasomatism of a moderately depleted mantle (Perinelli et al., 2008; Sapienza and Scribano, 2000; Sapienza et al.,
73
2005). The residual nature of the peridotites was also confirmed by
the depletion observed in HFSE relative to primordial mantle abundances (data not shown), similar to that reported by Sapienza and
Scribano (2000).
The chondrite-normalized REE pattern of pyroxenites is displayed
in Fig. 5, both for whole rocks and clinopyroxenes. The pyroxenite
samples show a REE upward-convex pattern, characterized by a
less-pronounced enrichment of the more incompatible elements
(i.e., La, Ce, and Pr), and of HREE compared to MREE. Among the analyzed samples, only the composite peridotite–pyroxenite sample,
XIC-26, show a different pattern, whose mineralogical composition
was somewhat transitional between that of peridotites and pyroxenites. The enrichment of LREE in pyroxenites relative to chondrite
varies among the studied samples (Lan/Ybn = 2.4–11.3), opposite to
what was observed in peridotites (Fig. 5a and b). Following
Sapienza and Scribano (2000), the pyroxenites represent the crystallization product of deep magmatic liquids that intruded the peridotites at different levels of the lithospheric mantle. In this framework,
the differences in LREE enrichments among pyroxenite samples
could reflect varying degrees of metasomatism, depending on the extent to which the metasomatizing melts interact with the surrounding peridotite.
Fig. 6 plots Zr/Nb ratios versus Zr concentrations of the bulk rocks.
The complete data set, comprising our data plus those in the literature (Sapienza and Scribano, 2000), define two clearly distinguishable compositional fields for pyroxenites and peridotites:
pyroxenites are characterized by Zr concentrations of 26–40 ppm
and a Zr/Nb ratio of ~20, while peridotites exhibit a much lower Zr
content of ~ 8 ppm and a Zr/Nb ratio of ~ 4. The lower Zr concentration
of peridotites is related to their more refractory nature. A particularly
notable behavior is displayed by the XIC-26 pyroxenite sample, which
has Zr and Zr/Nb values of 3.5 ppm and 5.3, respectively, which are
much more similar to those of peridotites. Fig. 6 also shows the compositions of erupted lavas having HIMU and DM signatures. Zr incompatibility makes lavas obviously richer in this element than xenoliths;
nevertheless, the Zr/Nb ratio is little affected by crystal-melt fractionation processes, so that lavas and xenoliths are directly comparable.
Whereas pyroxenites exhibit the Zr/Nb ratio that is typical of DM, peridotites clearly fall in the HIMU range.
Sr- and Nd-isotope data measured in clinopyroxenes from both
peridotites and pyroxenites are listed in Table 2 and plotted in
Fig. 7. The peridotites show almost homogeneous 87Sr/ 86Sr and
143
Nd/ 144Nd values of ~ 0.7029 and ~ 0.5129, respectively, while the
pyroxenites exhibit variable 87Sr/ 86Sr (0.7028–0.7031) and 143Nd/
144
Nd ≈ 0.5130. As shown in Fig. 7, our values are consistent with
those reported by Bianchini et al. (2010) and Tonarini et al. (1996).
Consideration of the complete data set indicates the absence of any
appreciable differences in the 87Sr/ 86Sr ratios among peridotites and
pyroxenites, whereas their 143Nd/ 144Nd ratios differed slightly. In
particular, the isotope ratios were slightly higher in the pyroxenites
(between 0.5129 and 0.5130) than in the peridotites (between
0.5128 and 0.5129), highlighting the presence of two distinct compositional groups. In accordance with inferences from the Zr/Nb ratio,
inspection of Fig. 7 also suggests that the pyroxenites formed by
melts coming from a deep mantle (probably DM-type) source that intruded into the shallower peridotite mantle level (with signatures
similar to a HIMU-type source).
6. Chemical and isotope compositions of noble gases from
fluid inclusions
As already noted, fluid inclusions occur inside olivine and pyroxene crystals of Hyblean ultramafic xenoliths. These fluid inclusions
represent a primary gaseous phase (dominated by CO2) coexisting
with growing minerals at mantle depths, as demonstrated by thermobarometric and microthermometric studies carried out in the same
74
A. Correale et al. / Lithos 132-133 (2012) 70–81
Table 1
Whole rock and mineral phases major element compositions of studied peridotites and pyroxenites xenoliths.
Peridotites
Sample
wt.%
SiO2
TiO2
Al2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P2O5
Cr2O3
XIH-1
WR
Opx
Cpx
WR
WR
Ol
Opx
Cpx
WR
Cpx
WR
Opx
Cpx
Spl
Grn
WR
Opx
Cpx
Spl
Grn
WR
Oliv
Opx
Cpx
Spl
WR
Oliv
Opx
Cpx
Spl
41.93
56.66
53.27
44.07
42.36
41.00
55.24
52.23
25.46
48.33
44.18
54.93
50.03
n.a
40.36
47.63
51.23
48.48
0.09
41.60
42.95
40.81
55.42
52.80
n.a
50.40
40.61
55.80
51.64
n.a
0.06
0.11
0.27
0.03
0.08
b.d.l
0.16
0.32
1.13
1.73
1.50
0.11
1.19
0.45
0.11
1.24
0.54
1.59
0.56
0.38
1.07
n.a
0.19
0.31
0.11
0.20
n.a
0.12
0.67
0.42
0.93
2.76
3.66
0.77
1.53
0.03
4.05
5.50
23.81
9.53
11.17
4.14
6.25
65.14
22.62
9.24
6.89
9.52
59.78
23.22
8.21
n.a
3.62
2.75
55.80
3.90
n.a
2.85
5.59
33.10
9.11
6.25
2.84
8.74
9.11
10.39
6.72
3.88
12.92
7.43
7.75
8.06
3.95
11.72
16.16
8.88
12.53
4.26
20.35
13.50
10.57
9.80
6.26
2.25
11.61
6.70
9.48
6.65
3.44
19.60
0.11
0.15
b.d.l
0.11
0.11
0.26
0.08
0.15
0.11
0.14
0.11
0.20
0.11
0.09
0.46
0.15
0.19
0.10
0.11
0.35
0.13
0.15
0.15
0.08
0.09
0.10
0.15
0.16
0.13
0.20
39.46
32.67
0.00
40.29
36.15
47.58
31.00
16.09
13.55
12.63
14.41
32.01
15.32
22.14
14.12
14.77
27.20
13.54
18.02
17.32
17.95
49.80
32.94
16.70
20.80
26.70
50.16
34.30
16.86
16.15
3.85
0.57
22.68
1.59
2.99
0.12
1.39
19.59
8.89
18.59
16.07
0.72
21.29
n.a
6.75
15.87
1.35
17.81
n.a
5.57
13.66
n.a
0.47
22.98
n.a
9.20
n.a
0.35
19.16
n.a
0.04
0.49
1.11
0.03
0.12
0.70
0.64
1.35
0.73
1.74
0.76
0.00
0.79
n.a
0.03
1.20
0.18
1.41
n.a
n.a
0.80
n.a
n.a
0.51
n.a
0.30
n.a
0.01
1.36
n.a
0.02
b.d.l
b.d.l
0.02
0.02
0.01
0.03
b.d.l
0.02
b.d.l
0.09
n.a
n.a
n.a
n.a
0.07
n.a
n.a
n.a
n.a
b.d.l.
n.a
n.a
n.a
n.a
b.d.l.
n.a
n.a
n.a
n.a
0.06
b.d.l
b.d.l
0.03
0.16
b.d.l
b.d.l
b.d.l
0.03
b.d.l
0.11
n.a
n.a
n.a
n.a
0.08
n.a
n.a
n.a
n.a
0.07
n.a
n.a
n.a
n.a
0.10
n.a
n.a
n.a
n.a
0.05
0.47
0.85
0.13
0.35
0.08
0.70
1.05
0.08
0.05
0.02
n.a.
n.a.
0.52
n.a.
0.03
n.a.
n.a.
0.11
n.a.
0.11
n.a.
0.35
0.65
11.40
0.68
n.a.
0.59
1.14
29.15
XIH-2
XIH-3
Pyroxenites
XIP-4
XIP-17
XIP-14
XIP-28
XIC-26
Table 2
Trace elements abundance of whole rock and mineral phases and Sr–Nd isotopic compositions of handpicked clinopyroxenes from studied Hyblean enclaves.
Peridotites
Sample
ppm
Ba
Sc
Sr
V
Ce
Co
Cs
Dy
Er
Eu
Gd
Hf
Ho
La
Lu
Nb
Nd
Pb
Pr
Rb
Sm
Ta
Tb
Th
Tm
U
Y
Yb
Zr
87
Sr/86Sr
143
Nd/144Nd
Pyroxenites
XIH-1
XIH-1
XIH-1
XIH-2
XIH-3
XIH-3
XIH-3
XIH-3
XIP-4
XIP-4
XIP-17
XIP-17
XIP-14
XIP-28
WR
OPX
CPX
WR
WR
OLIV
OPX
CPX
WR
CPX
WR
CPX
WR
WR
5.65
6.35
202
35.50
5.25
129
0.80
0.26
0.10
0.11
0.36
0.14
b.d.l.
3.05
0.02
2.95
1.90
0.70
0.58
1.15
0.40
0.07
0.05
0.35
b.d.l.
0.36
1.10
0.10
7.00
n.a.
n.a.
0.24
18.41
0.57
93.96
0.21
58.91
0.07
0.12
0.12
0.08
0.08
0.07
0.05
0.11
0.05
0.18
0.14
0.09
0.07
0.09
0.10
0.04
n.a.
0.07
0.05
0.08
0.57
0.18
0.94
n.a.
n.a.
0.43
94.86
276
272.99
41.22
25.57
0.01
1.91
0.88
0.83
2.40
0.61
0.34
20.11
0.11
1.97
15.35
0.93
4.17
0.04
2.73
0.22
n.a.
1.61
0.12
0.34
8.88
0.79
17.07
0.702956 ± 5
0.512917 ± 6
2.60
8.45
186
35.25
4.01
121.50
0.15
0.19
0.09
0.07
0.22
0.03
b.d.l.
2.15
0.01
1.00
1.35
b.d.l.
0.47
0.68
0.30
0.60
b.d.l.
0.33
b.d.l.
0.07
0.75
b.d.l.
1.88
n.a.
n.a.
4.90
8.75
216
60.50
4.69
123.50
0.10
0.31
0.15
0.14
0.43
0.16
0.06
2.50
0.02
1.90
2.20
0.70
0.60
0.60
0.50
0.05
0.06
0.20
b.d.l.
0.12
1.40
0.10
9.40
n.a.
n.a.
0.02
3.20
0.02
5.75
0.01
134.45
0.01
0.01
0.01
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
b.d.l.
0.01
0.01
0.01
0.00
0.02
0.01
b.d.l.
n.a.
b.d.l.
b.d.l.
0.01
0.05
0.02
0.11
n.a.
n.a.
0.07
21.22
2.27
117
0.36
71.49
0.01
0.19
0.13
0.05
0.14
0.09
0.04
0.09
0.03
0.24
0.34
0.02
0.06
0.03
0.11
0.02
n.a.
0.01
0.02
0.01
1.10
0.18
5.50
n.a.
n.a.
0.59
58.45
250
239
32.61
29.58
0.07
2.17
0.96
1.11
3.01
0.80
0.38
12.71
0.12
1.85
17.14
0.20
4.08
0.15
3.52
0.30
n.a.
0.36
0.14
0.16
9.51
0.81
46.29
0.7030315 ± 6
0.512919 ± 7
17.90
23.90
60.10
386
6.42
100
0.20
2.13
1.00
0.77
2.46
0.90
0.38
2.10
0.10
1.00
6.30
b.d.l.
1.26
0.45
2.20
0.05
0.33
0.10
0.12
0.08
8.00
0.60
26.20
n.a.
n.a.
0.44
55.70
94.61
332
12.46
41.82
0.02
4.27
2.00
1.43
4.63
2.28
0.78
3.18
0.22
0.59
13.00
0.06
2.26
0.08
4.11
0.12
n.a.
0.06
0.26
0.02
18.93
1.63
48.73
0.7031435 ± 5
0.512947 ± 8
181
42.20
648
401
18.00
n.a.
n.a.
n.a.
n.a.
1.26
n.a.
1.90
n.a.
7.80
0.18
n.a.
12.00
n.a.
n.a.
9.00
3.29
0.10
0.60
0.50
n.a.
0.20
18.00
1.20
45.00
n.a.
n.a.
1.51
57.87
57.69
380
7.47
31.04
0.02
2.85
0.93
1.18
3.90
2.11
0.44
1.71
0.06
0.10
9.79
0.25
1.52
0.12
3.58
0.02
n.a.
0.02
0.10
0.01
10.14
0.52
41.45
0.702859 ± 7
0.512994 ± 6
26.00
37.70
161
284
16.00
n.a.
n.a.
n.a.
n.a.
1.00
n.a.
1.30
n.a.
5.60
0.16
n.a.
11.00
n.a.
n.a.
9.00
2.87
n.a.
0.50
0.30
n.a.
0.10
17.00
1.16
40.00
n.a.
n.a.
3.20
n.a.
150
231
10.60
n.a.
n.a.
2.28
1.11
0.96
2.75
1.54
0.42
3.67
0.11
2.06
10.00
n.a.
1.74
0.99
2.79
0.18
0.41
0.23
0.13
0.08
11.70
0.78
39.30
n.a.
n.a.
A. Correale et al. / Lithos 132-133 (2012) 70–81
75
peridotites this work
pyroxenites this work
peridotites Sapienza & Scribano (2000)
pyroxenites Sapienza & Scribano (2000)
100
Zr/Nb
DM
10
HIMU
1
1
10
100
1000
Zr (ppm)
Fig. 6. Zr/Nb vs Zr diagram for whole rock from peridotites and pyroxenites. The Zr/Nb
ratio of peridotites approaches slightly those of a HIMU-type source whereas those DM
are more similar to a HIMU source.
Reference fields: DM source from Sun and McDonough (1989), Hofmann (1988), Arevalo and McDonough (2010) while HIMU source from Chauvel et al. (1992).
Fig. 4. C1-normalized REE patterns of a) whole rock (this work and Sapienza and Scribano, 2000) and b) cpx (this work and Perinelli et al., 2008) from peridotites.
Normalization to C1 is after Anders and Grevesse (1989).
a
100
rock/C1
10
1
XIP-4
XIP-28
XIP-17
XIC-26
XIP-14
area (Sapienza et al., 2005; Tonarini et al., 1996). Fluid inclusions also
retain noble gases, which can be used as powerful tracers of the mantle source.
The concentrations of noble gases in the studied xenoliths are
listed in Table 3. The He content varies from 7.3 × 10 − 14 to
2.6 × 10 − 11 mol/g in mineral separates of peridotite nodules, and
from 5.1 × 10 − 13 to 3.4 × 10 − 11 mol/g in those of pyroxenite (Fig. 8).
The He abundance in peridotites overlapped the range of data
reported for Hyblean samples by Sapienza et al. (2005). The Ar concentration was measured for the first time in both peridotites and pyroxenites, varying from 1.89 × 10 − 13 to 6.64 × 10 − 12 mol/g and from
8.2 × 10 − 13 to 3.2 × 10 − 11 mol/g, respectively. Samples that are richer
in He are generally also richer in Ar; this behavior was observed in all
of the investigated samples, although the He/Ar ratios did differ between the samples.
The He and Ar concentrations differ systematically among the
cogenetic minerals (olivines, orthopyroxenes, and clinopyroxenes)
of each sample (see Table 3), with them being slightly higher in clinopyroxenes and orthopyroxenes than in olivines. The 3He/ 4He ratios
were also the highest in clinopyroxenes and orthopyroxenes. This
partially agrees with the findings of Sapienza et al. (2005), who
reported generally comparable values in olivine and orthopyroxene
but lower values than in cogenetic clinopyroxenes.
It is well known that fluid inclusions can be contaminated by air.
In order to evaluate the air contribution in our samples, we plotted
4
He/ 20Ne versus 40Ar/ 36Ar ratios, as shown in Fig. 9. The 4He/ 20Ne
ratio varies between 61 and 4740, while that of 40Ar/ 36Ar varies
0.1
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
0.5132
DM
100
143
10
rock/C1
peridotites this work
pyroxenites this work
peridotites Bianchini et al. (2010); Tonarini et al. (1996)
pyroxenites Bianchini et al. (2010); Tonarini et al. (1996)
0.5131
Nd/144Nd
b
0.5130
0.5129
HIMU
HIMU
Cpx XIP-4
Cpx XIP-17
1
0.5128
Cpx XIC-26
0.5127
0.7026 0.7027 0.7028 0.7029 0.7030 0.7031 0.7032 0.7033
87
Sr/86Sr
0.1
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Fig. 5. C1-normalized REE patterns of a) whole rock from this work (symbols) and
Sapienza and Scribano (2000) (shaded area) and b) cpx (this work) from pyroxenites.
Normalization to C1 as in Fig. 4.
Fig. 7. Sr–Nd isotopic composition of Hyblean xenoliths from this work and from literature (Bianchini et al., 2010; Tonarini et al., 1996). Nd isotopes of peridotites fall fully in
the range of HIMU source, while higher ratio in pyroxenites may testify a certain tendency toward a DM end member.
The DM and HIMU values are from Zindler and Hart (1986).
76
A. Correale et al. / Lithos 132-133 (2012) 70–81
Table 3
Noble gases analyses of mineral phases from peridotitic and pyroxenitic Hyblean xenoliths.
Mineral
phase
Weight
(g)
4
He 10− 13
(mol/g)
Peridotites
XIH-1
Ol
XIH-1
Opx
XIH-1
Cpx
XIH-2
Ol
XIH-2
Opx
XIH-3
Ol
XIH-3
Opx
XIH-3
Cpx
3.0
1.7
2.2
1.4
1.6
1.0
3.2
3.3
2.0
1.6
1.7
0.7
2.0
1.9
2.1
1.2
1.1
0.7
1.2
0.7
4.2
3.4
4.2
4.3
2.1
2.1
10.0
5.9
11.4
10.6
66.6
45.4
168.6
102.7
263.0
188.5
5.6
0.4
0.7
0.5
0.3
0.4
0.3
0.3
2.8
6.5
3.7
2.2
3.3
2.2
7.5
4.2
5.5
5.1
Pyroxenites
XIH-4
Cpx
XIH-14
Cpx
XIC-26
Cpx
XIP-28
Cpx
XIH-17
Cpx
0.8
0.6
0.5
0.5
0.4
0.6
0.6
0.5
0.2
0.5
0.3
0.3
0.5
0.5
302.1
337.3
13.9
30.4
11.5
10.5
5.1
116.2
83.6
170.4
136.2
45.8
88.8
81.4
35.4
14.7
22.7
8.4
6.3
3.9
0.3
107.9
6.2
4.7
5.3
9.3
9.8
13.0
Sample
20
Ne 10− 15
(mol/g)
40
Ar 10− 12
(mol/g)
4
3
He/4He
(R/Ra)
Err +/−
40
Err
(%)
1.4
1.5
2.2
2.0
1.9
1.3
1.1
0.2
1.1
1.0
1.9
3.6
2.7
2.3
5.7
4.1
6.6
4.3
211.9
180.2
637.9
685.3
1236.5
1034.8
596.6
762.5
359.2
90.7
310.3
489.2
2033.0
2036.1
2245.8
2453.8
4739.9
3668.6
7.0
7.0
7.0
7.0
7.0
7.0
7.1
7.1
7.2
7.3
7.2
7.2
7.2
7.2
7.2
7.2
7.4
7.4
0.19
0.25
0.11
0.12
0.13
0.10
0.11
0.08
0.06
0.07
0.06
0.07
0.06
0.06
0.06
0.06
0.05
0.05
314.8
312.2
342.7
357.1
445.7
414.9
323.4
513.7
479.6
461.8
611.1
804.8
426.8
414.1
502.1
553.5
716.4
641.8
0.05
0.05
0.06
0.07
0.07
0.06
0.03
0.02
0.04
0.05
0.02
0.27
0.03
0.07
0.05
0.06
0.04
0.05
1.3
0.9
1.4
1.0
0.7
1.1
2.2
2.6
2.3
1.7
1.1
0.5
8.2
6.9
7.2
5.3
6.7
8.1
25.6
13.6
2.4
1.6
2.7
0.8
0.8
32.3
6.9
3.7
3.9
4.5
20.7
23.9
853.9
2294.1
61.4
360.8
183.0
269.7
1833.5
107.7
1342.8
3660.9
2580.4
490.3
904.1
624.8
0.3
7.3
7.3
7.6
7.6
7.4
7.2
7.1
7.4
7.3
7.4
7.4
7.4
7.4
7.3
0.06
0.05
0.10
0.10
0.11
0.10
0.13
0.10
0.05
0.06
0.06
0.08
0.05
0.05
620.1
900.9
315.9
338.3
314.9
342.7
550.4
323.3
649.2
554.2
477.1
431.4
938.6
556.6
295.5
0.10
0.04
0.03
0.07
0.42
0.07
0.04
0.07
0.27
0.04
0.04
0.04
0.02
0.04
2.3
3.7
8.8
15.4
6.8
9.0
1.3
4.2
2.2
9.8
9.1
3.2
0.6
0.7
Air
between 316 and 939. All of the samples plot close to a computed
curve of the binary mixing between an atmospheric term and a hypothetical MORB source (Graham, 2002; Marty et al., 1983), thereby
confirming atmospheric contamination of the gases released from
fluid inclusions (Fig. 9; see caption for further details). It is noteworthy that the highest 4He/ 20Ne and 40Ar/ 36Ar ratios (indicating our
samples with the lowest air contamination) were generally measured
in samples with the highest gas contents released from fluid inclusions. The most likely causes of the atmospheric signature are (1)
air contamination in the mantle due to subduction of atmospheric
components (e.g., Sarda, 2004) and (2) air entrapment in microcracks
of minerals during or after the eruptive activity (e.g., Nuccio et al.,
2008). Regarding the first cause, we recall that gas emissions at
He/20Ne
Ar/36Ar
4
He/40Ar*
Mofeta dei Palici – which is located in the northern Hyblean area
and close to Quaternary volcanic systems – showed 40Ar/ 36Ar values
in the range 1600–2000, which are consistent with a mixing between
air and a MORB mantle (Nakai et al., 1997; INGV-PA database). These
ratios are much higher than those measured in our fluid inclusions.
Given that the Hyblean mantle surely has 40Ar/ 36Ar ratios above
2000, the low 40Ar/ 36Ar ratios of fluid inclusions cannot be inherited
from the mantle but instead are probably caused by air contamination
that occurs at shallow levels or after the entrapment of fluid inclusions. Similar conclusions have been previously drawn by fluid inclusions studies from other areas (Graham, 2002; Martelli et al., 2011;
Nuccio et al., 2008; Porcelli and Ballentine, 2002). At the present
10000
8.0
He/20Ne
4
7.5
MORB
100
10
1
7.0
3
He/4He(R/Ra)
1000
peridotites
AIR
0.1
250
6.5
1.E-14
350
pyroxenites
450
550
650
750
850
950
40
Ar/36Ar
1.E-13
1.E-12
1.E-11
1.E-10
He (mol/g)
Fig. 8. 3He/4He (expressed as R/Ra) ratio vs He concentration in the investigated
samples.
Fig. 9. 4He/20Ne vs 40Ar/36Ar ratios of fluid inclusions from Hyblean xenoliths. The
curve defines a mixing trend between two end-members: 1) MORB, having 4He/20Ne
~ 10,000 (Marty et al., 1983), 40Ar/36Ar ~ 40,000 (Graham, 2002); 2) Air, having 4He/
20
Ne = 0.318, 40Ar/36Ar = 295.5.
A. Correale et al. / Lithos 132-133 (2012) 70–81
state of knowledge, we therefore believe that air components
entrapped in microcracks of minerals during or after their eruption
provide the most likely explanation of air contamination in our fluid
inclusions (Ballentine and Barfod, 2000).
We tested if air contamination could affect the 3He/ 4He ratios by
means of the formula of Giggenbach et al. (1993) that uses the 4He/
20
Ne ratio of the atmospheric end-member to evaluate the degree of
contamination of a sample. The results demonstrate that these corrections have practically negligible effects. The He contents in fluid
inclusions were in fact practically unmodified by air contamination
due to the low concentration of He (5.2 ppm) in air. For the same reason, even when considering a fractionated air (e.g. air saturated water
or air saturated sea water) as a contaminant, the 3He/ 4He values
would remain unchanged after correction. The 40Ar concentration in
fluid inclusions was corrected by assuming that all of the 36Ar found
in the samples was of atmospheric origin, according to the following
reported equation:
40
40
Ar ¼ Armeasured −
40
36
36
Ar= Ar Armeasured Þ
air
where 40Ar* represents the corrected 40Ar. This equation allowed us
to also compute a corrected-for-air 4He/ 40Ar ratio, hereafter referred
as 4He/ 40Ar* (see Table 3).
The 3He/ 4He values vary between 7.0 and 7.4 Ra in the peridotites,
in accordance with those observed by Sapienza et al. (2005), while
they vary between 7.2 and 7.6 Ra in the pyroxenites. The 4He/ 40Ar*
ratios range between 0.4 and 8 in the peridotites and between 0.6
and 15 in the pyroxenites, indicating partial overlap in the values.
The 3He/ 4He ratios are plotted versus the 4He/ 40Ar* ratios in Fig. 10.
In general, the average 3He/ 4He and 4He/ 40Ar* ratios were lower in
peridotite samples than in pyroxenite samples.
7. Discussion
7.1. Noble gases as geochemical tracers of mantle processes
In order to account for the observed variations of 3He/ 4He and
He/ 40Ar* ratios, we need to consider the main processes, both posteruptive and mantle-related, that can affect the noble-gas signature.
4
7.1.1. Post-eruptive processes
Post-eruptive processes that could affect the variability of 3He/ 4He
ratios in fluid inclusions are cosmogenic 3He production and
77
radiogenic 4He. Samples were collected from surfaces in rapid erosions, and the XIH1 and XIH3 samples were from road cuts, which
should have made massive ingrowth of cosmogenic helium highly
unlikely. In principle, the crushing procedure should release only
gas retained in the bubbles and not matrix-sited components such
as post-eruptive 3He and 4He, further preventing both cosmogenic
and radiogenic contributions implanted in the crystal matrix. Also,
our data on peridotites display 3He/ 4He values (7.0–7.4 Ra) that overlap those of Sapienza et al. (2005; 7.0–7.6 Ra) for similar samples, despite us using a single-step crushing while Sapienza et al. (2005) used
prolonged crushing (strokes for 2.5 min). It is reasonable to assume
that if the samples were rich in a post-magmatic component that
could be released by crushing, very different crushing techniques
should give different results.
In addition, considering that the two principal parameters that
control the post-eruptive production of 3He and 4He (age of the sample and exposure at the surface) are similar for pyroxenites and peridotites, and that for the same He concentration the 3He/ 4He ratio is in
most cases higher for pyroxenites than for peridotites (Fig. 8), we attribute this isotopic difference to genetic processes rather than to
post-eruptive processes. Therefore, even if we cannot definitely exclude slightly alteration of the original 3He/ 4He ratio of individual
samples by post-eruptive processes, the mean difference between peridotites and pyroxenites should be largely attributable to mantle
processes.
7.1.2. Mantle processes
In their study of worldwide mantle xenoliths, Yamamoto et al.
(2009) observed that 4He/ 40Ar* decreased from a typical mantle
value of 1–4 (Graham, 2002; Ozima and Podosek, 1983) down to
0.1, paralleled by a decrease in 3He/ 4He from 7 to 3 Ra. They attributed this to kinetic fractionation among noble-gas atomic species due to
their diffusion through the mineral assembly of mantle, towards
magma channels crossing the mantle itself. Because of their high incompatibilities, noble gases would be preferentially partitioned in
the magma relative to mantle minerals so as to diffuse from mantle
to magma flowing through the channels. Under such conditions, the
different diffusivities of 3He, 4He, and 40Ar would induce a kinetic
fractionation of these isotopes, and so the 3He/ 4He and 4He/ 40Ar* ratios would decrease in the mantle source. Following the approach of
Yamamoto et al. (2009), we calculated how the noble-gas ratios
should vary as a result of the diffusive fractionation (see Yamamoto
et al., 2009, for further details on boundary conditions). The process produces a dramatic decrease in 3He/ 4He compared to that
8.0
DM
Magma degassing
7.6
7.4
7.2
3
4
He/ He (R/Ra)
7.8
peridotites this work
pyroxenites this work
mixing curve
diffusive fractionation
7.0
6.8
HIMU
6.6
0.1
1.0
10.0
4
3
4
4
40
100.0
40
He/ Ar*
Fig. 10. Plot of He/ He (R/Ra) corrected for air contamination vs He/ Ar* ratios of peridotites and pyroxenites. The dashed curves result from a mixing, which is a consequence of
metasomatic processes, between a DM and a HIMU end-member. The dark thick arrow indicates the variations of 4He/40Ar* ratio during degassing processes. See the text for more
information about DM (light gray area) and HIMU (deep gray area) sources. The meaning of the diffusive fractionation curve (thin curve) is exposed in the text (Section 7.1.2).
A. Correale et al. / Lithos 132-133 (2012) 70–81
1.0
vesicles
0.1
0
200
400
600
800
1000
1200
Pressure (MPa)
Fig. 11. Variations of 4He/40Ar* ratio in melt (gray curves) and vesicles (black curve)
during closed (dashed curve) and open (continuous curve) system degassing due to
the decompressive ascent of magma. The shaded areas show the variation range of
the 4He/40Ar* ratio for the vesicles during a hypothetical magmatic depressurization
from ~900 to 700 MPa (see the text for details). The equilibrium degassing model
and volatile solubilities were the same as reported by Paonita and Martelli (2007) for
a basalt melt at 1200 °C. The initial conditions were H2O and CO2 contents of 0.3 wt.%
and 1 wt.%, respectively, and an initial He/Ar* ratio of 3, in accordance with a pristine
basaltic magma from the upper mantle (Paonita and Martelli, 2007).
of 4He/ 40Ar*, showing a conflicting behavior with respect to the trend
observed in the data set (Fig. 10).
In order to explain the 4He/ 40Ar* variations observed in our dataset we focused on the process of magma degassing, given that the different solubility of noble gases in silicate melts could generate large
changes of 4He/ 40Ar* (e.g., Burnard, 2004; Moreira and Sarda, 2000;
Paonita and Martelli, 2006, 2007, and references therein). Indeed as
magma ascends throughout the mantle as a consequence of depressurization, the noble gases leave the magma in proportion to their
solubilities. Specifically, due to the solubility of He being higher
than that of Ar (e.g., Iacono-Marziano et al., 2010; Nuccio and
Paonita, 2000), the degassing process would increase the 4He/ 40Ar*
ratio of the residual magma. Based on the equilibrium degassing calculations of Paonita and Martelli (2007) for a typical basalt coming
from the upper mantle and exsolving CO2-dominated fluids, the first
vesicles that separate from a melt with 4He/ 40Ar ≈ 3 have a 4He/
40
Ar ratio of ~ 0.2 (see Fig. 11). This ratio would increase as the
magma ascends. Fluid inclusions entrapped at different pressures in
the forming minerals can hence record 4He/ 40Ar* ratios reflecting
variable extents of degassing. By assuming an open-system degassing,
a hypothetical pressure decrease from ~900 to 700 MPa – which is
within the range of the expected depths of the investigated products
(27–35 km; Sapienza et al., 2005) – is readily able to explain the observed 4He/ 40Ar* variations (Fig. 11).
In accordance with Paonita and Martelli (2007), the kinetic fractionation of 3He and 4He during magma degassing of CO2-dominated
fluids (like our fluid inclusions) can be excluded, and hence a further
process must occur in conjunction with degassing to explain the measured 3He/ 4He variations. Starting from the mentioned petrologic and
geochemical evidences, we propose that the He-isotope variability results from a mixing between heterogeneous mantle sources, where
two local end-members having different 3He/ 4He values can be identified: (1) the peridotite domain having 3He/ 4He ≈ 7 Ra and (2) the
pyroxenite domain with 3He/ 4He ≈ 7.6 Ra.
In this regard it is important to note that the values of peridotites
are closer to those of a HIMU-type mantle source ( 3He/ 4He = 5–7 Ra;
Hanyu and Kaneoka, 1998; Moreira and Kurz, 2001), while those of
pyroxenites approach those of a DM-type source ( 3He/ 4He = 8 ± 1
Ra; Allègre et al., 1995). Accordingly, the deep high- 3He/ 4He pyroxenite melt decompresses during ascent, reaching the low- 3He/ 4He
a
7.5
80%
60%
40%
20%
7.4
He/4He (R/Ra)
4
He/40Ar
melt
7.3
7.2
7.1
3
10.0
peridotitic levels. Open-system degassing processes can easily increase the 4He/ 40Ar ratios of the pyroxenite from values starting at
around 0.4, such as for an early vapor separated from melt having
typical mantle ratio (see above), up to 15 or even more (Fig. 11).
The subsequent mixing process would occur between the high- 3He/
4
He pyroxenite having variable 4He/ 40Ar ratios and a low- 3He/ 4He
peridotite mantle with 4He/ 40Ar = 0.4–1.0 (Fig. 10). This range,
which is slightly lower than the typical mantle ratios of 1–4 (Ozima
and Podosek, 1983), could result from the extraction of liquids produced by partial melting of the primordial peridotite. In fact, due to
the incompatibility being higher for He than for Ar
(DHe = 1.17 × 10 − 4 versus DAr = 1.10 × 10 − 3; Heber et al., 2007),
~1% of melting would account for the required decrease in He/Ar in
the residual peridotite. Fig. 10 sketches the process using a set of
two end-member mixing curves between a fixed term and other
end-members with different 4He/ 40Ar ratios. The grid clearly shows
that the described process may easily explain the complete data set.
In a previous study of Hyblean peridotitic xenoliths, Sapienza et al.
(2005) had already observed 3He/ 4He values in the range 7.3 ± 0.3 Ra,
although they did not investigate pyroxenitic xenoliths, and they proposed a deep metasomatizing source. Based on our results, the
Hyblean pyroxenites would therefore represent a metasomatizing
agent (having 3He/ 4He ≈ 7.6 Ra) that is located in a deeper portion
of the local mantle, while the peridotites would correspond to a shallower layer (having 3He/ 4He ≈ 7 Ra) that is occasionally crossed by
pyroxenite melts ascending from depth (Perinelli et al., 2008;
Sapienza and Scribano, 2000; Scribano et al., 2008). This would result
in the shallower portion of the Hyblean mantle being partially or totally refertilized by such metasomatizing melts.
It is noteworthy that the measured 3He/ 4He values of the Hyblean
xenoliths (7.0–7.6 Ra) were slightly but distinguishably higher than
7.0
Pyroxenites
Peridotites
6.9
0
2
4
6
8
10
12
14
Nd (ppm)
b
7.5
60%
40%
7.4
He/4He (R/Ra)
100.0
80%
20%
7.3
7.2
7.1
3
78
7.0
Peridotites
Pyroxenites
6.9
0
0.5
1
1.5
2
2.5
3
3.5
4
Sm (ppm)
Fig. 12. Plot of 3He/4He (R/Ra) vs Sm (a) and Nd (b) concentrations of whole rock from
Hyblean peridotites and pyroxenites. The curves describe a mixing trend between two
hypothetical end-members associated respectively with peridotitic and pyroxenitic
sources. Ticks indicate percentages of the pyroxenitic end-member. The He concentrations of the two end-members, used to calculate the mixing path, were the highest
measured in pyroxenites and the lowest in peridotites. Accordingly, the ratio between
the He contents of these two end-members (which determines the convexity of the
curve) is ~ 460.
A. Correale et al. / Lithos 132-133 (2012) 70–81
those measured in all of the other mantle xenoliths from Europe and
North Africa (5.6–7.0 Ra; Beccaluva et al., 2007, 2008; Dunai and Baur,
1995; Gautheron et al., 2005; Martelli et al., 2011). Also, such values
are the highest measured in recent basaltic lavas of Italian volcanism
(Martelli et al., 2008, and references therein; Marty et al., 1994), with
only free gases from Pantelleria Island reaching similar values
(Parello et al., 2000).
7.2. Evidence of mantle metasomatism by coupling noble-gas and traceelement data
The geochemical investigations of Hyblean xenoliths carried out in
this study have suggested the existence of a vertical heterogeneity of
Hyblean mantle, with it being characterized by a pyroxenitic deep
layer and a peridotitic shallow portion that occasionally is partially
or totally metasomatized. We subsequently therefore investigated
the quantitative relationships between trace elements and 3He/ 4He
data.
Fig. 12 shows average 3He/ 4He versus Sm and Nd concentrations
of single nodules. Pyroxenites and peridotites defined two distinct
end-members of Hyblean mantle, where the highest 3He/ 4He data
were found in the pyroxenites that also showed the highest Sm and
Nd concentrations, while the same correspondence at the lowest
values was observed in the peridotites. The two mantle layers, characterized by extreme Sm, Nd, and 3He/ 4He values of the complete
data set, would be (1) the peridotitic one, with an average 3He/
4
He ≈ 7 Ra, Sm ≈ 0.25 ppm, and Nd ≈ 0.1 ppm, and (2) the pyroxenitic one, characterized by an average 3He/ 4He ≈ 7.4 Ra, Sm ≈ 3.5 ppm,
and Nd ≈ 13 ppm.
Different degrees of metasomatism were modeled by a hypothetical mixing between the two end-members, as already observed for
the noble gases and the REE data independently. Fig. 12 shows that
the strongly convex shape of the resulting mixing curves is in good
agreement with the data, suggesting that the metasomatic process
controls both trace-element and noble-gas geochemistry. Whereas
Sm and Nd mix linearly, the higher 3He/ 4He ratios in pyroxenites,
coupled to the high He content, makes He extremely sensitive to
metasomatic events. In fact, a very low contribution of fluids from pyroxenite dramatically changes the isotope ratio of peridotite mantle.
The scarcity of both He and incompatible trace elements accounts
for a strongly depleted character of Hyblean peridotites, which probably also suffered extensive degassing during melt extraction.
The He-isotope ratio gives some indications about the widely debated genesis of pyroxenites on a worldwide scale, with the two
main groups of interpretations being (e.g., Bodinier and Godard,
2003; Downes, 2007) (1) crystal precipitation from deep-mantle
magmas in conduits passing through the lithosphere and (2) recycling and recrystallization of subducted components belonging to
old oceanic crust in convecting mantle. Our data indicate the high
He-isotope ratio of Hyblean pyroxenites, which is close to the mean
value of MORBs, intended as samples of the convecting upper mantle
(Graham, 2002). Our data are thus consistent with the Hyblean pyroxenites originating from the first of the two hypothesized mechanisms, while processes connected to recycled components by
subducted crust should involve a lower He-isotope ratio due to
time-integrated 4He production from U and Th radiogenic decay.
8. Inferences from the heterogeneity of the Hyblean mantle
The geochemistry of Zr/Nb, Sr and Nd isotopes, and noble gases in
our mantle xenoliths highlights that Hyblean peridotites and pyroxenites resemble well-known HIMU and DM mantle sources, respectively. Furthermore, the petrographic evidence (peridotitic nodules
veined from pyroxenites) account for a deeper origin for the pyroxenites, so that the mantle layer having HIMU characteristics would be
shallower than the DM-type one. Such inferences raise two main
79
questions: (1) can a recycled component be associated with a peridotite mantle and a DM be associated with a pyroxenite mantle, and (2)
what is the meaning of their peculiar vertical stratification?
With regard to the first question, Sobolev et al. (2008) showed
that the enriched component is frequently linked to reaction pyroxenite, whereas the depleted component is likely to be derived from a
peridotitic source, which would contradict our results. However, by
studying mantle xenoliths from the Canary Islands, Gurenko et al.
(2009) suggested that a recycled mantle component is not necessarily
linked to an eclogite–pyroxenite paragenesis. In fact, it could also
exist in the form of peridotite when an old (>1 Ga) recycled component had sufficient time to be stirred back into the peridotite matrix.
The results of our study are thus consistent with the second hypothesis, and suggest that the mineralogical assemblage of recycled mantle component can range from pyroxenites, “hybrid melts”, up to
peridotites.
Concerning the inferred vertical stratification, our results contrast
with the view that the Euro-Mediterranean HIMU is generated by upwelling of a deep plume (e.g., Hoernle et al., 1995). In fact, when linking the HIMU signature to a mantle plume of recycled material, we
would expect the HIMU to originate from deeper than the MORB.
However, different models support our inferred location of HIMU
being shallower than MORB.
Scribano et al. (2008) assumed the presence of a serpentinitehosted hydrothermal system in the Hyblean lithosphere as a result
of tectonic uncovering and seafloor exposure of the uppermost mantle since middle Triassic. This hypothesis could account for the HIMU
marker of peridotites through the hydrothermal addition of U to altered rocks, especially by serpentine formation (Michard and
Albarède, 1985). The homogeneous isotope marker of He in olivines,
orthopyroxenes, and clinopyroxenes of the peridotite samples
would support the radiogenic production of 4He from serpentine
veins and its uniform diffusion toward the three mineralogical
phases. Since the serpentine veins were mainly found in microcracks
within olivine or along grain boundaries, we expect the 3He/ 4He ratios to be more radiogenic in olivines than in pyroxenes. This would
also require migration of He from the lattice into fluid inclusions, although we showed that this process seems to be of minor importance
(see Section 7.1.1). The more primitive He-isotope composition of pyroxenites would also imply either that the latter ones infiltrated the
peridotite matrix subsequent to the hydrothermal circulation causing
the HIMU signature or that the analyzed pyroxenites sampled portions of pyroxenite veins that were physically distant from the contact area with the surrounding peridotite, so as to avoid the main
interaction with serpentinizing fluids.
Indeed, the most striking feature of the investigated HIMU signature is its widespread occurrence in several Euro-Mediterranean
areas (Cebria and Wilson, 1995; Macera et al., 2003; Wilson and
Bianchini, 1999; Wilson and Downes, 1992), so that any explanation
should preferentially involve a regional scale. Lustrino and Wilson
(2007) summarized some of the different models that support the
presence of a Euro-Mediterranean HIMU mantle without invoking a
mantle plume actively upwelling from a thermal boundary layer at
the core–mantle boundary. The signature of recycled component
may be simply inherited by the shallow upper mantle in the form of
a metasomatized lithosphere or enriched asthenosphere, as suggested by the SUMA model (statistical upper mantle assemblage)
(Meibom and Anderson, 2004). Based on this, Piromallo et al.
(2008) explained the common HIMU-like character of erupted lavas
in different tectonic environments of the Euro-Mediterranean region
by proposing a sublithospheric dragging of the plume head located
in the Canary–Cape Verde zone as a result of a north-to-eastward migration of the Eurasian and African plates away from the hot spot.
This would have allowed a spreading of plume material in the shallow
sublithospheric mantle so as to produce a geochemically anomalous
(HIMU-like) level located above the depleted upper mantle. We
80
A. Correale et al. / Lithos 132-133 (2012) 70–81
therefore conclude that the somewhat anomalous mantle stratification that we have suggested here for Hyblean mantle can be
explained within the framework of the above model.
9. Conclusions
A comprehensive geochemical study of mantle-derived ultramafic
xenoliths hosted in some Hyblean volcanic systems was performed to
better characterize the lithospheric mantle below this area, thereby
expanding the existing knowledge. The investigated nodules consist of
peridotites and pyroxenites entrapped in some Miocene-age diatremes.
New data of major and trace elements coupled to Sr and Nd isotopes are reported here and compared to those available from similar
studies. The present study integrates these geochemical data with an
investigation of noble gases of fluid inclusions hosted in the same xenoliths paragenesis.
The obtained data led us to the following conclusions:
1) The mantle below the Hyblean area is heterogeneous, featuring a
shallower peridotitic layer with more evolved geochemical characteristics ( 3He/ 4He ≈ 7 Ra, 143Nd/ 144Nd ≈ 0.5129, and Zr/
Nb ≈ 4) relative to a deeper pyroxenite domain that shows a primitive character ( 3He/ 4He ≈ 7.6 Ra, 143Nd/ 144Nd ≈ 0.5130, and Zr/
Nb ≈ 30). Peridotites and pyroxenites seem to display HIMU and
DM affinities, respectively.
2) Metasomatic processes occur in the lithosphere below this area.
Particularly, deep pyroxenite melts were identified as a metasomatizing agent. By ascending toward the surface, they intrude
the peridotite mantle at different levels by partially or totally
refertilizing it.
3) The metasomatic processes control both trace-element and noblegas geochemistry. In contrast, previous studies of mantle xenoliths
(i.e., Matsumoto et al., 2000, and references therein) found a
decoupling between noble gases in fluid inclusions and radiogenic
isotopes and trace elements in the whole rock can very often be
observed in the same samples. The present study has revealed
that different geochemical tracers can display very different sensitivities to the effects of metasomatic mixing between two endmembers, and hence this process should be carefully considered
when formulating hypotheses of the processes underlying decoupling between noble-gas and trace-element geochemistry.
Acknowledgments
We thank Ilenia Arienzo for performing Sr and Nd isotope analyses
at INGV-Osservatorio Vesuviano and Mariano Tantillo for help in
noble-gas analyses. Sivio Rotolo is also thanked for assistance during
SEM-EDS analyses. This work is part of the PhD thesis of A.C. financially
supported by the Università di Palermo. Comments by two anonymous
reviewers and Editor-in-Chief A. Kerr greatly improved the article.
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