Redox state of subcontinental lithospheric mantle and relationships with metasomatism: insights from...

ORIGINAL PAPER

Redox state of subcontinental lithospheric mantleand relationships with metasomatism: insights from spinelperidotites from northern Victoria Land (Antarctica)

Cristina Perinelli • Giovanni B. Andreozzi •

Aida M. Conte • Roberta Oberti • Pietro Armienti

Received: 8 January 2012 / Accepted: 28 July 2012 / Published online: 28 August 2012

� Springer-Verlag 2012

Abstract Rift-related Cenozoic alkaline mafic lavas

from northern Victoria Land (Antarctica) carry abundant

mantle xenoliths whose oxygen fugacities (fO2) were

determined to assess how the metasomatism, related to

Cenozoic magmatism, affected the state of oxidation of

the lithospheric mantle. The xenoliths used for this study

are anhydrous spinel peridotites sampled in two localities,

Greene Point and Baker Rocks, that show different

extents of metasomatism: these are limited to incompati-

ble element enrichments in Greene Point and to enrich-

ments in major, minor and trace elements at Baker Rocks.

The data set includes a composite xenolith from Baker

Rocks, formed by a depleted lherzolite crosscut by an

amphibole-bearing vein. Mossbauer spectroscopy was

used to accurately determine the Fe3?/Fetot ratios in spinel

and amphibole minerals. Amphiboles were also charac-

terized by Single-Crystal X-ray Diffraction, and the

crystallographic data were used to calculate the dehy-

drogenation. The oxidation state recorded by the xenoliths

ranges from 0.2 to 1.5 log-bar units below the fayalite–

magnetite–quartz (FMQ) buffer (DlogfO2) with the high-

est values observed in the metasomatized samples from

Greene Point. For the vein of composite Baker Rocks

xenolith, Dlog fO2 was estimated on the basis of the

amphibole in -1.7 log-bar units, a value close to those

calculated for all Baker Rocks xenoliths (DlogfO2 = -1.5

to -1.1 log-bar units). These results indicate a similar

oxidation state for lithospheric mantle prior to the meta-

somatic event at Greene Point and Baker Rocks

(DlogfO2 * -1.3 log-bar units). Metasomatism produced

different effects in the shallow mantle at the two sites. At

Greene Point, an oxidizing metasomatic melt caused the

rise of fO2 in peridotite portions close to melt conduits up

to FMQ. In contrast, at Baker Rocks, a metasomatizing

melt with fO2 similar to that of the peridotite matrix

produced chemical changes in the surrounding mantle

rocks and amphibole crystallization without significantly

modifying the local oxidation state. The origin of such

different metasomatic melts, as well as the variety of

primary magmas produced during the magmatic phase of

Cenozoic rifting, is linked to the geodynamic evolution

of the rift and probably involved the melting of a heter-

ogeneous mantle source composed of a peridotite veined

by pyroxene-bearing veins formed from an earlier

amagmatic phase of the rift.

Keywords Upper mantle oxygen fugacity � Mantle

metasomatism � Rift evolution � Northern Victoria Land �Antarctica

Communicated by C. Ballhaus.

C. Perinelli (&) � P. Armienti

Dipartimento di Scienze della Terra, Universita degli Studi di

Pisa, Via S. Maria 53, 56126 Pisa, Italy

e-mail: [email protected]

P. Armienti


G. B. Andreozzi

Dipartimento di Scienze della Terra, Sapienza Universita di

Roma, P.le A. Moro 5, 00185 Rome, Italy


A. M. Conte

Istituto di Geoscienze e Georisorse (IGG-CNR) UOS Roma,

P.le A. Moro 5, 00185 Rome, Italy


R. Oberti

Istituto di Geoscienze e Georisorse (IGG-CNR) UOS Pavia,

via Ferrata 1, 27100 Pavia, Italy


123

Contrib Mineral Petrol (2012) 164:1053–1067

DOI 10.1007/s00410-012-0788-7

Introduction

Oxygen fugacity (fO2) is a key factor in controlling magma

genesis and metasomatism in the mantle. In fact, mantle

fO2 constrains mineral assemblages and the speciation of

C–O–H fluids (CO, CO2, CH4, H2 and H2O), which in turn

play a primary role in determining the pressure (P) and

temperature (T) conditions for the onset of partial melting.

Many studies that focused on upper mantle fO2 have shown

the presence of vertical and lateral heterogeneities, which

were ascribed to partial melting (Bryndzia and Wood 1990;

Woodland et al. 2006), tectonic setting (Wood 1990;

Ballhaus 1993; Parkinson and Arculus 1999) and metaso-

matism (McGuire et al. 1991; Ballhaus 1993; McCammon

et al. 2001).

Spinel peridotite xenoliths included in Cenozoic volca-

nic rocks of northern Victoria Land (NVL), Antarctica, are

used to assess the redox state and its possible evolution

during Cenozoic magmatism, in this part of lithospheric

mantle.

Cenozoic magmatism in the NVL is linked to the West

Antarctic Rift System in a region of consistent crustal

thinning (the Moho in the coastal area is at *20–25 km of

depth; Trehu et al. 1989) as revealed by a topographical

trough running from the Antarctic Peninsula to the Ross

Embayment–northern Victoria Land (Fig. 1) (LeMasurier

and Thomson 1990; Tessensohn and Worner 1991; Beh-

rendt et al. 1991). Many lines of evidences point to the

origin of Cenozoic igneous activity in NVL consistent with

a passive rifting frame related to a transtensive tectonic

regime; this is also in agreement with the geothermobaro-

metric estimates on xenoliths from lower crust and upper

mantle (Worner 1999; Armienti and Perinelli 2010).

NVL mantle xenoliths revealed heterogeneities in the

mantle composition, which were interpreted in terms of

partial melting episodes and metasomatism (Coltorti et al.

2004, Perinelli et al. 2006; Melchiorre et al. 2011). In

particular, samples hosted by magmas from two localities

only 80 km apart show distinct metasomatic features. In

Baker Rocks samples (BR: Mt. Melbourne Volcanic

Province), metasomatism is evident by the occurrence of

amphibole (both disperse and in vein) as well as by the

enrichment in major, minor (Fe and Ti, respectively) and

incompatible trace elements. On the contrary, in Greene

Point (GP) samples, the metasomatism is cryptic and is

only recorded by clinopyroxene through selective

Meander Intrusives McMurdo Volcanics Ultramafic Xenoliths

This work

Other outcrops

Weddell Sea

Transantarctic M

ountains

Ross Sea

90 W90 E

Inferred boundary of the West Antarctic Rift System

60S

0

Cenozoic dike swarms

Fig. 1 Sketch map of northern

Victoria Land and sampling

localities. McMurdo Volcanics

enclose the Cenozoic volcanic

products (Kyle 1990), while

Meander Intrusives enclose the

intrusive-subvolcanic varieties

(Tonarini et al. 1997)

1054 Contrib Mineral Petrol (2012) 164:1053–1067

123

enrichment in incompatible elements. Metasomatic chan-

ges on BR and GP lithospheric mantle have been inter-

preted as results of ‘‘wall-rock’’ metasomatism and/or

‘‘diffuse’’ metasomatism: the first is linked to the transport

of melt in fractures (veins and dykes); the latter is related to

percolation of small melt fractions along grain boundaries

in the peridotite matrix. The shift from ‘‘wall-rock’’ to

‘‘diffuse’’ metasomatism is controlled by the different

ratios of melt/rock interaction volume involved during the

metasomatism (Xu and Bodinier 2004 and reference

therein). Mineralogical and chemical modifications of BR

xenoliths were indeed induced by both ‘‘wall-rock’’

(presence of amphibole and/or Fe–Ti enrichments) and

‘‘diffuse’’ metasomatism (variable enrichments in incom-

patible trace elements in clinopyroxene), while only dif-

fusive processes acted at Greene Point.

Perinelli et al. (2006) related the metasomatic event to

Cenozoic magmatism and documented a similar metaso-

matic agent for both suites, even if its oxidation state was

not evaluated. Our working hypothesis is that mineralogi-

cal and geochemical differences observed in BR and GP

spinel peridotites contain information on the oxidation state

of the metasomatic melts. To verify this hypothesis,

information on fO2 conditions was retrieved from a selec-

tion of variably metasomatized BR and GP peridotite

xenoliths and used to decipher the complex evolution of

lithospheric mantle in northern Victoria Land.

Materials and methods

Petrographic and mineralogical features of spinel

peridotites

Mantle xenoliths selected for this study (Table 1) occur in

a melanephelinite (Greene Point) and in an alkali basalt

(Baker Rocks) belonging to the McMurdo Volcanic Group,

the Unit that collects the volcanic products of Cenozoic

igneous activity related to the rift of Ross Sea (Kyle 1990).

Details on petrography and mineral chemistry of GP and

BR samples are reported in Perinelli et al. (2006) and

Armienti and Perinelli (2010). The main textural and

compositional features of GP and BR samples are briefly

described here. Xenoliths from both localities consist of

coarse-textured spinel lherzolites (12–6 % of clinopyrox-

ene) and spinel harzburgites (5–3 % of clinopyroxene).

Table 1 Spinel peridotite xenoliths from lithospheric mantle of northern Victoria Land (NVL) selected for the present study

Sample Rock

type

T (�C) P (GPa) Cr# in

Spinel

F % Sp-based

(Hellebrand

et al. 2001)

F % Cpx-based

(Hellebrand

et al. 2002)

Shape of Cpx

REE pattern

(La/Yb)N

in CpxcType of

metasomatism

GP6 Hz 1032 1.74 0.39 15 15b Spoon LREE-

enriched

1.6 Cryptic

GP22 Hz 1048 1.1 0.47 17 15 Convex-upward

LREE-enriched

22.4 Cryptic

GP31 Hz 1109 1.45 0.24 10 7b Flat LREE-

enriched

0.7 Cryptic

GP40 Lh 985 1.53 0.20 8 9b Normal LREE-

depleted

0.01 Unmetasomatized

GP42 Lh 1113 1.64 0.11 2 3 Normal LREE-

depleted

0.04 Unmetasomatized

BR213 Lh 907 1.1 0.27 11 – Convex-upward

LREE-enriched

2.8 Cryptic

BR218 Lh 911 0.9 0.19 8 10b Spoon LREE-

enriched

4.2 Cryptic

BR218* Lh 927a 0.9 0.22 9 – Convex-upward

LREE-enriched

4.6 Fe–Ti

BR219 Hz 933 1.21 0.16 6 9b Spoon LREE-

enriched

2.8 Cryptic

Temperatures are estimated by the geothermometer of Ballhaus et al. (1991). Pressure data are from Armienti and Perinelli (2010)

GP greene point, BR baker rocks, Hz harzburgite, Lh lherzolite, Cr# Cr number expressed as Cr/(Cr ? Al) atom, F % estimated degree of partial

melting, Sp spinel, Cpx clinopyroxene

* Lherzolite portion adjacent to amphibole-bearing vein (see text)a Temperature estimated using the composition of olivine-spinel pair adjacent to amphibole-bearing veinb Partial melting occurred partly in the garnet field stability (Perinelli et al. 2006)c (La/Yb)N ratio in clinopyroxene is normalized with respect to C1-Chondrite (McDonough and Sun 1995). REE data are from Perinelli et al.

(2006)

Contrib Mineral Petrol (2012) 164:1053–1067 1055

123

They are generally homogeneous, with the only exception

of a sample from Baker Rocks (BR 218, Fig. 2), where

the lherzolite is crosscut by an amphibole-bearing vein

*2-mm wide. Minerals forming the peridotite xenoliths

exhibit in general uniform composition: the forsterite (Fo)

component of olivine ranges from 89 (lherzolite) to 92

(harzburgite); orthopyroxenes and clinopyroxenes can be

referred to as enstatite (En88–91, Wo2, Fs7–10) and

Cr-diopside (Wo44–48, En46–52, Fs4–6), respectively, and

their compositional range is consistent with the degree of

partial melting suffered by the source (e.g., the Al2O3

content in clinopyroxene decreases from 6.88 to 3.19 wt %

in GP42 lherzolite and in GP22 harzburgite, respectively).

Both ortho- and clinopyroxenes show selective enrichment

in incompatible trace elements, which have been related to

a metasomatic event (Perinelli et al. 2006). The chromium

number (Cr# = Cr/(Cr ? Al) ratio) of Cr-rich spinels

increases from 0.11 in lherzolites to 0.47 in harzburgites, as

expected in a mantle affected by variable degrees of partial

melting. In the sample BR218, all phases have higher Fe

contents than in the xenoliths with similar lithology. In

particular, mineral phases of the peridotite portion closest

to the amphibole-bearing vein (hereafter named BR218*,

see Table 1) are enriched in iron; spinel is depleted in

Al2O3 and enriched in Cr2O3, and clinopyroxene and spinel

are also enriched in titanium (Table 2).

In the vein, amphibole occurs as follows: (1) small

euhedral grains (B0.1 mm in diameter), in direct contact

with the peridotite matrix, classified as Ti–rich pargasite

according to Hawthorne and Oberti (2007) (average

#mg = 87, TiO2 = 3.70 wt % and Al2O3 = 14.65 wt %);

(2) larger crystals (up to *0.3 mm in diameter) of kaer-

sutitic composition (#mg = 85, TiO2 = 4.87 wt % and

Al2O3 = 13.51 wt % on average) occasionally surrounded

by a pale yellow glass with microcrystals of

clinopyroxene ? olivine ± plagioclase. At a distance of

2.5 cm far from the vein, only clinopyroxene still reveals

metasomatic features such as the selective enrichment of

incompatible trace elements. The gradual chemical evolu-

tion detected in this sample reflects the action of a me-

tasomatizing chromatographic process which occurred at

the scale of a mantle xenolith (Perinelli et al. 2006).

Mossbauer spectroscopy for Fe3?/Fetot determination

in spinel and amphibole

Eight mantle xenoliths (5 from Greene Point and 3 from

Baker Rocks) representative of unmetasomatized (GP40

and GP42) and differently metasomatized samples were

selected on the basis of texture (coarse type; Table 1) and

size, in order to provide enough material for mineral sep-

arates with minimal contamination by the host basalt. All

samples were crushed, and fresh spinels were picked up by

hand under a binocular microscope.

Since the amphibole-bearing vein that crosscuts the

BR218 peridotite can be reasonably considered as the

metasomatizing agent of this portion of upper mantle,

Mossbauer spectroscopy (MS) was also carried out on

amphibole itself in order to measure its Fe3?/Fetot ratio,

and hence allow calculation of fH2, and fO2 conditions in

the vein. The limited amount of available sample allowed

only Mossbauer analysis of the kaersutite-type amphibole

(hereafter named BR218Amph).

Mossbauer absorbers were prepared by pressing finely

ground samples, mixed with powdered acrylic resin to self-

supporting discs. The amount used corresponded to about

2 mg Fe/cm2, to avoid thickness effects. Spectra were

collected at Sapienza University of Rome, using a con-

ventional spectrometer system operating at room temper-

ature (298 K) in constant acceleration mode with a 57Co

source in rhodium matrix. Spectral data for the velocity

range -4 to ?4 mm/s were recorded in a multichannel

analyzer using 512 channels. After velocity calibration

against a high-purity a-iron foil spectrum, the raw data

were folded in 256 channels. The spectra are made of

doublets and were fitted using a five-doublet model by the

Recoil 1.04 (Lagarec and Rancourt 1998) fitting program

and assuming symmetrical Lorentzian peak shapes (Fig. 3).

The best fits were evaluated by reduced v2, and uncer-

tainties were calculated using the covariance matrix. Errors

were estimated to be close to ±0.02 mm/s for isomer shift

(d), quadrupole splitting (DEQ) and peak width (C). In the

present study, spectral areas for Fe2? and Fe3? were cor-

rected for the temperature effect with f factors calculated

by De Grave and Van Alboom (1991) for room tempera-

ture conditions (f2 = 0.687 and f3 = 0.887). Errors asso-

ciated with doublet areas are estimated to be less than

±3 % (Table 3).

Fig. 2 Plane-polarized light image of composite peridotite xenolith

BR218. Enlarged image shows the large kaersutitic amphiboles

crystallized in the vein that crosscuts the BR218 lherzolite


123

Table 2 Representative analyses of minerals of Greene Point and Baker Rocks spinel peridotite xenoliths

GP6 GP22 GP31

Ol Opx Cpx Sp Ol Opx Cpx Sp Ol Opx Cpx Sp

SiO2 40.68 56.32 53.45 0.04 40.57 56.33 53.02 0.02 41.02 54.74 52.67 0.04

TiO2 0.04 b.d.l. b.d.l. 0.03 0.04 0.01 0.05 0.10 b.d.l. 0.07 0.16 0.08

Al2O3 0.01 3.16 3.33 36.14 0.34 2.77 3.19 30.63 0.02 4.19 5.57 47.70

Cr2O3 n.a. 0.70 1.07 35.09 n.a. 0.73 1.39 41.12 n.a. 0.68 1.40 22.00

FeOtot 8.35 4.61 2.14 11.13 7.80 5.04 2.14 11.98 8.96 5.33 2.40 9.94

MnO 0.10 0.14 0.10 b.d.l. 0.05 0.14 0.05 b.d.l. 0.12 0.06 0.05 b.d.l.

NiO 0.25 n.a. n.a. 0.20 0.32 n.a. n.a. 0.16 0.34 n.a. n.a. 0.30

MgO 49.69 34.82 18.44 17.88 50.02 34.75 17.69 16.67 49.40 32.91 16.22 19.11

CaO 0.08 1.06 21.88 n.a. 0.10 1.01 20.77 n.a. 0.07 0.93 20.47 n.a.

Na2O n.a. b.d.l. 0.43 n.a. n.a. 0.09 0.99 n.a. n.a. 0.09 1.32 n.a.

K2O n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

Total 99.20 100.79 100.83 100.52 99.24 100.86 99.28 100.67 99.93 99.00 100.24 99.18

Fo 91.4 92.0 90.9

Mg# 0.93 0.94 0.74 0.92 0.94 0.71 0.92 0.92 0.77

Cr# 0.39 0.47 0.24

GP40 GP42 BR213

Ol Opx Cpx Sp Ol Opx Cpx Sp Ol Opx Cpxa Sp

SiO2 40.75 54.86 52.75 0.22 40.86 54.75 51.75 0.04 40.54 55.36 52.51 0.04

TiO2 0.01 0.02 0.18 0.09 0.02 0.08 0.41 0.14 b.d.l. 0.11 0.16 0.11

Al2O3 b.d.l. 4.04 4.89 50.04 0.02 4.08 6.88 57.92 b.d.l. 3.41 5.21 44.01

Cr2O3 n.a. 0.46 0.93 18.26 n.a. 0.53 1.01 10.46 n.a. 0.50 1.30 24.24

FeOtot 9.53 6.02 2.49 10.82 10.66 6.11 2.97 10.46 9.34 5.66 2.55 12.14

MnO 0.17 0.10 0.05 b.d.l. b.d.l. 0.17 0.02 b.d.l. 0.12 0.08 b.d.l. b.d.l.

NiO 0.42 n.a. n.a. 0.36 0.34 n.a. n.a. 0.35 0.38 n.a. n.a. 0.23

MgO 49.08 32.99 16.15 19.65 48.42 33.19 15.50 20.75 48.65 32.66 16.24 18.25

CaO 0.07 0.66 22.17 n.a. 0.09 0.74 20.20 n.a. 0.07 0.88 20.24 n.a.

Na2O n.a. 0.09 0.86 n.a. n.a. 0.06 1.42 n.a. n.a. 0.37 1.66 n.a.

K2O n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

Total 100.02 99.25 100.46 99.43 100.42 99.71 100.14 100.10 99.10 99.03 99.87 99.00

Fo 90.2 89.3 90.3

Mg# 0.91 0.92 0.76 0.91 0.90 0.78 0.91 0.92 0.73

Cr# 0.20 0.11 0.27

BR218 BR218* BR219

Ola Opxa Cpxa Sp Ol Opx Cpx Sp Pargasitic Ampha Kaersutitic Amph Ol Opx Cpx Sp

SiO2 40.24 55.38 52.39 0.03 39.84 55.05 51.33 0.03 42.72 42.27 40.69 54.35 52.48 0.04

TiO2 b.d.l. b.d.l. 0.49 0.13 0.00 0.12 0.87 0.51 3.66 4.87 b.d.l. 0.03 0.17 0.20

Al2O3 b.d.l. 4.00 5.19 50.27 0.01 4.55 6.07 46.91 14.82 13.51 0.01 4.23 5.09 52.22

Cr2O3 n.a. 0.50 0.88 17.93 n.a. 0.46 0.90 19.69 0.82 0.90 n.a. 0.54 0.94 15.28

FeOtot 11.97 6.70 2.84 14.28 14.53 8.22 3.74 15.53 4.62 5.15 9.39 6.75 2.64 11.42

MnO 0.02 b.d.l. 0.18 b.d.l. 0.23 0.19 0.11 b.d.l. b.d.l. 0.05 0.18 0.11 0.05 b.d.l.

NiO 0.35 n.a. n.a. 0.29 0.30 n.a. n.a. 0.29 n.a. n.a. 0.38 n.a. n.a. 0.35

MgO 47.31 33.10 15.54 17.35 45.25 31.38 15.68 16.68 16.38 15.87 48.87 32.77 16.11 19.44

CaO 0.07 0.76 21.46 n.a. 0.07 0.90 21.16 n.a. 11.68 11.51 0.08 0.83 22.23 n.a.

Na2O n.a. b.d.l. 1.31 n.a. n.a. 0.09 0.94 n.a. 2.98 2.81 n.a. 0.06 0.78 n.a.

K2O n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. 0.81 0.91 n.a. n.a. n.a. n.a.

Total 99.96 100.44 100.28 100.28 100.22 100.97 100.80 99.64 98.49 97.84 99.59 99.68 100.49 98.93

Fo 87.6 84.7 90.3

Mg# 0.90 0.91 0.68 0.87 0.88 0.66 0.86 0.85 0.90 0.92 0.75

Cr# 0.19 0.22 0.16

Fo forsterite component in olivine; Mg# [MgO/(MgO ? FeOtot) molar]; Cr# [Cr3?/(Cr3??Al3?) atomic]; n.a. not analyzed; b.d.l. below detection limit

* Minerals close to amphibole-bearing veina Data from Perinelli et al. (2006)


123

Crystallographic and crystal-chemical analysis

of BR218Amph.

Single-crystal structure refinement (SREF) and crystal-

chemical analyses were done in order to understand in

detail the crystal-chemical mechanisms ruling dehydroge-

nation and to evaluate the O2- component in BR218

amphibole. The data collection was done with a Philips

PW1100 single-crystal 4-circle diffractometer equipped

with graphite monochromated Mo-Ka X radiation. The

unit-cell parameters were calculated from least-squares

refinement of the d* values obtained for 60 rows of the

reciprocal lattice by measuring the center of gravity of each

reflection and of its antireflection in the 2h range -70 to

70�. Two monoclinic equivalents were collected in the 2h

range 4–70�, and corrections were applied for absorption

and Lorentz polarization. Reflections with Io [ 3 r(I) were

considered as observed during unweighted full-matrix

least-squares refinement on F done using a program spe-

cifically written at CNR-IGG Pavia to deal with complex

solid-solutions. Scattering curves for fully ionized scatter-

ing species were used at sites where chemical substitutions

occur; neutral versus ionized scattering curves were used at

the T and anion sites. Atom coordinates, equivalent iso-

tropic displacement parameters and their anisotropic

components are reported in Table 4.

Site populations were calculated from the unit formula

and validated based on the structure refinement and the

methods developed for amphiboles at CNR-IGG Pavia

(Oberti et al. 2007). In particular, the populations of the

M(1), M(2) and M(3) sites were derived from refined site-

scattering values and mean bond lengths, and those of the

T(1) and T(2) sites based on the mean bond lengths and

selected compositional parameters. The O2- component

was estimated based on its relation with the M(1)–M(2)

distance, which had been calibrated by SIMS analysis of a

suite of partially dehydrogenated pargasite–hastingsite–

kaersutite (R2 = 0.99). The crystal was then mounted in

epoxy, polished and analyzed by electron microprobe

(same procedures as for the thin section). The crystal-

chemical formula, obtained combining electron microprobe

analyses (EMPA), structure refinement and Mossbauer

results, is discussed in a section below.

Analytical results

Quantification of Fe3?/Fetot ratios in spinel

and amphibole

Fe3?/Fetot ratios measured by Mossbauer analyses (Fe3þMS)

in spinel grains are reported in Table 3. The same table

also reports Fe3þ=Fetot ratios calculated from EMP analy-

ses on basis of charge balance by assuming a perfect

stoichiometry for the mineral (Fe3þEMPA). Fe3þ=Fetot ratios

determined by MS on GP spinels range from 0.17 to 0.35;

in contrast, those determined on BR spinels are lower and

more homogeneous, ranging from 0.13 to 0.15. Notably,

for GP spinels, Fe3þ=Fetot ratios show a rough correlation

with metasomatism, whereas for BR spinels, there is no

significant difference between cryptically and modally

metasomatized samples (Tables 1, 3).

Indeed, for all samples, Fe3?/Fetot ratios derived by

Mossbauer analysis are systematically higher than those

calculated from EMPA assuming spinel stoichiometry, the

largest differences being observed for metasomatized GP

xenoliths (Table 3; Fig. 4). It is well known that the

expcalc.

Fe3+

Fe2+

velocity (mm/s)

tran

smit

tanc

e (%

)

-4 0-1-2-3 1 2 3 4

99.2

99.4

99.6

99.8

100.0

99.0

Fig. 3 Typical 57Fe Mossbauer spectrum of spinels from NVL

mantle xenoliths, collected at room temperature

Table 3 Fe3?/Fetot ratios and oxidation degree in spinels of NVL

mantle xenoliths

Sample Fe3?/Fetot from

Mossbauer

Fe3?/Fetot from

EMPA

z (%)

GP6 0.35 0.06 26

GP22 0.25 0.04 19

GP31 0.31 0.01 27

GP40 0.17 0.13 5

GP42 0.22 0.11 9

BR213 0.15 0.12 1

BR218 0.14 0.06 6

BR218* 0.13 0.10 2

BR219 0.14 0.12 1

BR218Amph 0.24

* Close to amphibole vein; BR218Amph: kaersutite-type amphibole

within amphibole-bearing vein in BR218; errors on Fe3?/Fetot ratios

estimated at maximum ±0.03; z %: oxidation degree of spinels cal-

culated after Quintiliani et al. (2006)


123

assumption of spinel stoichiometry may produce large

uncertainties on Fe3?/Fetot ratios quantification due to the

following: (1) systematic errors associated with the very

low amounts of magnetite component in mantle spinels

(commonly \ 10 mol %; Canil et al. 1990); (2) underes-

timation of Fe3? content in the spinel arising from possible

non-stoichiometry linked to Fe3? excess and cation

vacancies (Andreozzi and Lucchesi 2002; Bosi et al. 2004;

Lenaz et al. 2004; Quintiliani et al. 2006).

As pointed out by Bosi et al. (2004), the Fe3? excess

(the difference between Fe3þMS and Fe3þ

EMPA) can reflect

secondary oxidation processes. In particular, Fe3? of pri-

mary mineral formation (Fe3þP ) must be distinguished from

Fe3? excess derived by secondary oxidation. In this pro-

cess, the oxidation of primary Fe2þ Fe2þP

� �is matched by

the formation of cation vacancies (h). Menegazzo et al.

(1997) and Quintiliani et al. (2006) used crystal-chemical

data of spinel samples to calculate their degree of oxidation

z ¼ 8h= 3�hð ÞFe2þP . The parameter z may be used to

reconstruct the oxidation history of spinel samples: absence

of secondary oxidation corresponds to z = 0 %, full oxi-

dation corresponds to z = 100 %. Moreover, at z = 0 %,

the spinel composition is perfectly stoichiometric, and

therefore, Fe2þP and Fe3þ

P contents may be effectively rep-

resented by Fe2? and Fe3? calculated from EMPA data.

Values of z obtained for GP and BR spinels range from 5

to 27 % and from 0 to 6 %, respectively (Table 3; Fig. 5).

On conservative basis, only z values higher than 10 %

should be used to infer oxidation event(s). Therefore, GP

metasomatized spinels are characterized by minor to

medium oxidation, whereas BR spinels do not show any

clear evidence of oxidation. It is important to point out that

none of the crystals shows any evidence of alteration by

Table 4 Atom coordinates and atomic-displacement parameters (Beq, A2; bii 9 104) for the refined amphibole from amphibole-bearing vein in

BR218 xenolith (no. 1215 in the CNR-IGG-Pv database)

Atom x/a y/b z/c Beq b11 b22 b33 b12 b13 b 23

O(1) 0.10745 0.08658 0.21882 0.94 0.00226 0.00081 0.00866 -0.00017 0.00089 -0.00007

O(2) 0.11878 0.17171 0.72855 0.90 0.00218 0.00078 0.00826 -0.00004 0.00069 0.00026

O(3) 0.10793 0.00000 0.71446 1.09 0.00247 0.00094 0.01023 0.00000 0.00066 0.00000

O(4) 0.36580 0.25019 0.78686 1.19 0.00392 0.00078 0.01073 -0.00040 0.00146 0.00023

O(5) 0.35004 0.13961 0.10891 1.21 0.00270 0.00116 0.01010 -0.00007 0.00067 0.00115

O(6) 0.34610 0.11676 0.60813 1.25 0.00278 0.00113 0.01188 0.00007 0.00131 -0.00104

O(7) 0.34291 0.00000 0.27861 1.46 0.00312 0.00102 0.01716 0.00000 0.00099 0.00000

T(1) 0.28253 0.08523 0.30243 0.68 0.00188 0.00052 0.00632 -0.00008 0.00067 -0.00004

T(2) 0.29070 0.17271 0.81009 0.66 0.00175 0.00056 0.00580 -0.00011 0.00082 0.00005

M(1) 0.00000 0.08501 0.50000 1.16 0.00245 0.00147 0.00674 0.00000 0.00143 0.00000

M(2) 0.00000 0.17683 0.00000 0.69 0.00195 0.00057 0.00611 0.00000 0.00095 0.00000

M(3) 0.00000 0.00000 0.00000 0.81 0.00247 0.00057 0.00658 0.00000 0.00022 0.00000

M(4) 0.00000 0.27845 0.50000 1.00 0.00317 0.00074 0.01020 0.00000 0.00305 0.00000

A 0.00000 0.50000 0.00000 2.71 0.01071 0.00110 0.04164 0.00000 0.01957 0.00000

A(m) 0.05518 0.50000 0.10767 2.75 0.00941 0.00183 0.02934 0.00000 0.01000 0.00000

A(2) 0.00000 0.46926 0.00000 2.86 0.01034 0.00085 0.04949 0.00000 0.01934 0.00000

M(40) 0.00000 0.25438 0.50000 1.10

oxidation

0.0

0.1

0.2

0.3

0.4

0.0 0.1 0.2 0.3 0.4

Fe3+/Fetot EMPA

Fe3+

/Fe to

t Mss

baue

r

GP

BR

BR218*

1:1 lin

e

Fig. 4 Fe3?/Fetot ratios calculated from EMPA by assuming a perfect

stoichiometry for the spinels compared with Fe3?/Fetot ratios based

on Mossbauer Fe3? determination. The gray field about the 1:1 line

represents the confidence band for which spinels may be considered

not have been affected by secondary oxidation (z \ 10 %, see text).

GP Greene Point, BR Baker Rocks, BR218* portion of the lherzolite

adjacent to amphibole-bearing vein


123

weathering (atmospheric or superficial alteration); hence,

the observed oxidation must have occurred in the mantle.

The Fe3?/Fetot ratio measured on the amphibole

BR218Amph is 0.24 (Table 3). This value falls within the

range 0.2–0.5 typical for Ti–rich amphiboles in the upper

mantle (Popp and Bryndzia 1992; Canil and O’Neill 1996).

Cation distribution of the amphibole BR218Amph

and dehydrogenation

Crystal data and structure refinement results of relevance

are summarized in Table 5; the unit formula based on all

the available information (EMP, SREF and MS) is reported

in atoms per formula unit (apfu) in Table 6. Note the close

agreement between SREF and EMP analyses (126.25 vs.

126.45 electrons per formula unit, epfu, for all the cation

sites). Site populations for the T and M sites were calcu-

lated based on refined site-scattering values (ss, epfu) and

mean bond lengths (mbl, A). Tetrahedral mbl confirm that

all TAl is ordered at the T(1) site; the absence of TAl dis-

order usually indicates temperature of crystallization

B850 �C. For the M(1, 3) sites, the crystal-chemical

information to be taken into account is as follows: (1) a

M(1) - M(2) distance corresponding to 0.94 O2- apfu at

O(3) (Oberti et al. 2007); (2) the presence of significant Ti

(because the Beq value at M(1) is much higher than those at

the M(2) and M(3) sites) and of very low Fe3? content

(because the value of DM(1) is very low) at the M(1) site.

Taking into account these constraints and the refined ss and

mbl values, the best cation distribution we have obtained

(in apfu) is as follows:

M 1ð Þ : Mg1:33Fe2þ0:18Fe3þ

0:05Ti0:44 ssobs¼ 31:52;sscal¼ 31:62½ �;M 2ð Þ : Mg1:19Fe2þ

0:03Al0:48Fe3þ0:10Cr0:11Ti0:09

ssobs¼ 28:22;sscal¼ 28:45½ �;M 3ð Þ : Mg0:78Fe2þ

0:22 ssobs¼ 15:10;sscal¼ 15:08½ �:

Hence, the BR218Amph is kaersutite according to the

nomenclature of Leake et al. (1997), but will become

partially dehydrogenated Ti-rich pargasite in the new

nomenclature scheme for amphiboles recently approved

by IMA. What is more relevant to the present study,

the oxo component in BR218Amph is balanced almost

uniquely by the occurrence of Ti at the M(1) site

[Mð1ÞTi1Oð3ÞO2�

2Mð1Þ Fe2þ;Mg� �

�1Oð3ÞOH��2], a feature

that testifies crystallization in a low aH2O environment. In

contrast, the presence of M(1)Fe3? in BR218 kaersutite,

which would have testified post-crystallization dehydro-

genation according to the reaction Mð1;3ÞFe3þ1

Oð3ÞO2�1

Mð1;3ÞFe2þ�1

Oð3ÞOH��1, is very low (max 0.05 apfu) and

implies that the H loss in kaersutite during the xenolith

ascent, if any, is negligible. It follows that the oxo

-1.8

-1.6

-1.4

-1.2

-1.0

-0.8

-0.6

-0.4

-0.2

0.0

0 5 10 15 20 25 30

GP 31 F ~ 10%

GP 6 F ~ 15%

degree of oxidation z (%)

Δlog

fO2

GP 22 F ~ 15%

metasomatic oxidationno oxidation

BR218*close to amphibole-bearing vein

Δ log fO2 error bar

GP 42 F ~ 2%

GP 40 F ~ 8%

BR218 F ~ 8%

BR219F ~ 6%

BR213F ~ 11%

GP cryptically metasomatized

GPunmetasomatized

BR cryptically metasomatized

Fig. 5 DlogfO2 (fO2 relative to FMQ) for GP and BR mantle

xenoliths plotted against the oxidation degree parameter z. F is the

degree of partial melting for the two indicated xenoliths estimated by

the Hellebrand et al. (2001) method

Table 5 Crystal data and miscellaneous refinement information for kaersutite from amphibole-bearing vein in BR218 xenolith (no. 1215 in the

CNR-IGG-Pv database)

a (A) 9.8543(4) Crystal size (lm) 200 9 220 9 350 ss M(1) 31.57 T(1)–O (A) 1.672(1)

b (A) 18.0431(8) Radiation/monochromator MoKa/Graphite ss M(2) 28.22 T(2)–O (A) 1.638(1)

c (A) 5.3037(2) No. unique reflections 1380 ss M(3) 15.10 M(1)–O (A) 2.070(1)

b (�) 105.201(1) No. observed reflections (I C 3rI) 1270 ss M(4) 37.20 M(2)–O (A) 2.064(1)

V (A3) 910.02 Rmerge % 1.80 ss M(40) 1.94 M(3)–O (A) 2.065(1)

Sp. Gr. C2/m Robs % 2.26 ss A 3.34 M(1)–M(2) (A) 3.127(1)

Z 2 Rall % 2.44 ss A(m) 5.65 DM(1) 4.41

Dcalc (g cm-3) 3.17 Largest diff. peak (eA-3) 0.88 ss A(2) 3.48

ss refined site-scattering values (electrons per formula unit), DM(1) calculated according to Brown and Shannon (1973)


123

component (O2-*1 apfu) can be entirely ascribed to

crystallization conditions.

Redox conditions estimates

Oxygen fugacity calculations from spinel

One of the approaches most used to quantify the oxidation

state of shallow spinel-facies peridotite mantle is based on

the equilibrium reaction:

6Fe2SiO4 olivineð Þ þ O2 ¼ 3Fe2Si2O6 orthopyroxeneð Þþ 2Fe3O4 spinelð Þ ð1Þ

where Fe2SiO4, Fe2Si2O6 and Fe3O4 are the fayalite, fer-

rosilite and magnetite component in olivine (ol), orthopy-

roxene (opx) and spinel (sp), respectively.

Oxygen fugacity of BR and GP spinel peridotites was

estimated by two of the most accepted formulations, those

proposed by Wood (1990) and by Ballhaus et al. (1991).

The computed fO2 are expressed in log-bar unit relative to

the fayalite–magnetite–quartz (FMQ) buffer and are called

Dlog fO2 (Table 7). To calculate oxygen fugacity, we used

the Fe3?/Fetot value determined in spinel by Mossbauer

analyses, along with the equilibrium T and P values

reported in Table 1, with the exception for BR218*, for

which the equilibrium temperature was fixed at 900 �C

instead of 927 �C (Table 1) to take into account the

information obtained from SREF on amphiboles

(T B 850 �C) and on spinel (T = 909 �C, Perinelli et al., in

preparation). The use of T = 900 �C for calculation pro-

duces an uncertainty on Dlog fO2 (\0.1 log-bar units) well

below to that associated with the two methods used to

calculate fO2 (Dlog fO2 ± 0.5 log-bar units, Wood 1990;

Canil et al. 1990; Ballhaus et al. 1991).

Finally, the T–P uncertainties associated with each

investigated sample provide a maximum error on Dlog fO2

of ± 0.12 log-bar units.

The methods of Wood (1990) and Ballhaus et al. (1991)

yielded similar estimations for all samples at the level of

the calibration (differences are Dlog fO2 B 0.5 log-bar

units). Since both the Ballhaus et al. (1991) and Wood

(1990) models yield similar fO2 estimates and identify

similar variations of the local oxidation state, in the fol-

lowing discussion we chose to adopt the values calculated

following the first model.

Oxygen fugacity calculations from amphibole

The hydrogen fugacity, (fH2), and hence the oxidative

conditions, fO2, at which kaersutite BR218Amph crystal-

lized can be evaluated by the method proposed by Popp

et al. (1995; 2006). The Popp et al. (1995) approach is

Table 6 Chemical composition and unit formula (based on 24

anions) for kaersutite from amphibole-bearing vein in BR218 xenolith

wt % apfu

SiO2 42.27 Si 6.162

TiO2 4.87 Al 1.838

Al2O3 13.51 Sum T 8.000

Cr2O3 0.90 Ti4? 0.534

FeO* 3.92 Al 0.484

Fe2O3* 1.36 Fe3? 0.149

MnO 0.05 Fe2? 0.428

MgO 15.87 Mg 3.301

CaO 11.51 Cr 0.104

Na2O 2.81 Sum C 5.000

K2O 0.91 Ca 1.798

H2O** 1.00 Mg 0.145

F 0.19 Fe2? 0.050

Cl 0.03 Mn2? 0.007

–O=F -0.08 Sum B 2.000

–O=Cl -0.01 K 0.169

Total 99.12 Na 0.793

Sum A 0.962

ssEMPA 126.25 OH 0.972

ssSREF 126.45 O 0.933

F 0.088

Cl 0.007

Sum W 2.000

* FeO:Fe2O3 ratio calculated from SREF and Mossbauer results

** Calculated based on 24 (O, OH, F, Cl) with (OH ? F?Cl) = 2

apfu

Table 7 Oxygen fugacities for NVL mantle xenoliths

Sample DlogfO2

(Ballhaus

et al. 1991)

DlogfO2

(Wood

1990)

DlogfO2

from

BR218Amph

GP6 -0.23 -0.47

GP22 -0.54 -0.68

GP31 -0.65 -0.14

GP40 -1.35 -1.04

GP42 -1.20 -0.95

BR213 -1.21 -0.88

BR218 -1.14 -1.21

BR218* -1.52 -1.48 -1.71

BR219 -1.27 -0.89

DlogfO2: oxygen fugacity relative to FMQ buffer. In the first and

second columns, fO2 is based on spinel-bearing inter-crystalline

equilibrium (log-bar units, estimated error ±0.5); in the third column,

fO2 is based on amphibole intra-crystalline equilibrium (log-bar units,

estimated error ±1)

* Close to amphibole-bearing vein


123

based on the reaction expressed in terms of amphibole Ca,

Fe end-member:

Ca2Fe2þ5 Si8O22 OHð Þ2¼ Ca2Fe2þ

3 Fe3þ2 Si8O24 þ H2 ð2Þ

For this reaction, the equilibrium constant (K) is

expressed as

K ¼ f H2ð28:94Þ XFe3þð Þ2 Xvð Þ2

XFe2þð Þ2 XOHð Þ2U ¼ KXU ð3Þ

where v is the oxo component (i.e., the amount of O2-

instead of OH- at the O(3) site), U is the activity

coefficient term and KX is the thermodynamic mole

fraction term (i.e., the K expressed as mole fractions

rather than as activities). Popp et al. (2006) also expressed

the variation of KX as a function of T, P and amphibole

composition, with the equation:

log KX ¼ 4:23� 4380

T Kð Þ þ 2:61 Ti apfuð Þ � 0:42½ �f g

þ 88

T Kð Þ P kbarð Þ � 1½ ��

ð4Þ

which takes into account also the Ti content. The method is

calibrated experimentally, and its uncertainty is ±0.4 log-

bar units (Popp et al. 2006). In the case of BR218Amph,

using the crystal-chemical data, the measured Fe3?/Fetot

ratio by Mossbauer analyses and the P and T of 0.9 GPa

and 900 �C, respectively, the method provided estimates of

log fH2 = ?1.62 log-bar units.

As stated previously, the approach used by Popp et al.

provides the fH2 conditions of amphibole crystallization;

thus, the corresponding fO2 is to be determined using the

following equilibrium reaction:

H2O ¼ 1

2O2 þ H2 ð5Þ

for which the equilibrium constant (KH2O) is as follows:

KH2O ¼f H2 f O2ð Þ1=2

f H2Oð6Þ

In the environment in which mantle amphiboles formed,

the pressure of fluid is likely less than of the total pressure

(Pfluid \ Ptotal), and hence, the fH2O is a function of water

activity, aH2O:

f H2O ¼ f �H2O � aH2O ð7Þ

where f*H2O is the fugacity of pure H2O at a given P and T,

and aH2O (at the same P and T) is to be determined. Albeit

not commonly applied to mantle rocks, a method to cal-

culate water activity is the use of H2O-buffering equilibria

among the end-member components of peridotite minerals

(olivine, orthopyroxene, clinopyroxene, spinel/garnet and

amphibole; see ‘‘Dehydration equilibria’’ section in Popp

et al. 2006).

Following the criteria reported in Popp et al. (2006), and

keeping in mind the uncertainties involved in this approach

(Popp et al. 2006; Lamb and Popp 2009), we determined

the aH2O for BR218Amph by fixing the location of uni-

variant H2O-buffering equilibria as a function of P, T and

aH2O using the THERMOCALC software (Holland and

Powell 1990, 1998). Starting from the composition of the

co-existing minerals in BR218* (olivine, ortho- and

clinopyroxene and spinel plus amphibole, Table 2), the

values of aH2O were calculated for the following reactions

at 900 �C and 0.9 GPa.

2trþ 2fo ¼ 4diþ 5enþ 2H2O ð8Þ2trþ 2sp ¼ 2diþ 2catsþ 5enþ 2H2O ð9Þ2trþ 2sp ¼ 4diþ 3enþ 2mgtsþ 2H2O ð10Þ2trþ 3catsþ 2sp ¼ 7diþ 5mgtsþ 2H2O ð11Þ2trþ 2mgtsþ 2sp ¼ 4catsþ 7enþ 2H2O ð12Þ5trþ 2pargþ 7cats ¼ glþ 21diþ 9mgtsþ 6H2O ð13Þ5trþ 2parg ¼ glþ 14diþ 7enþ 2mgtsþ 6H2O ð14Þ7trþ glþ 12sp ¼ 2pargþ 10catsþ 21enþ 6H2O ð15Þ

where tr = tremolite, fo = fosterite, di = diopside,

en = enstatite, sp = spinel (MgAl2O4), cats = Ca-tscher-

mak, mgts = Mg-tschermakite, parg = pargasite and

gl = glaucophane are the relevant mineral end-members.

The aH2O values calculated for the dehydration equi-

libria (8)–(15) vary from a 0.02 (reaction 12) to 0.11

(reaction 8), except the reaction (11) for which the calcu-

lated aH2O is 0.25. The mean value of aH2O is 0.06 and

the1r of 0.02.

This estimate is consistent with the crystal-chemical

constraints (see above) that imply amphibole crystalliza-

tion in a low aH2O environment.

On the basis of aH2O (0.06) and KH2O (at appropriate

P and T; Bandura and Livov 2006), the fO2 conditions of

amphibole crystallization, and hence of the metasomatizing

agent can be estimated by the Eqs. (6) and (7) and result in

a value of Dlog fO2 = -1.71 log-bar units. The minimum

uncertainty on this value (±1 log-bar units) is mainly due

to the uncertainties involved in the method used to estimate

aH2O and fH2.

Discussion

Oxygen fugacity in northern Victoria Land lithospheric

mantle

Calculated fO2 values cluster below the FMQ buffer, and

Dlog fO2 varies from -1.52 to -0.23 log-bar units

(Table 7) well within the range accepted for subcontinental


123

lithospheric mantle (Dlog fO2 from -2.0 to ?1.0 log-bar

units; Ionov and Wood 1992; Ballhaus 1993; Woodland

et al. 2006).

No sample lacking metasomatic overprint has been yet

recovered at Baker Rocks; however, it is reasonable to

consider the less cryptically metasomatized sample

(BR219) as the best constraint to pre-metasomatic fO2.

Under this assumption, the lithospheric mantle underlying

both localities has a similar pre-metasomatic oxidation

state as testified by the close agreement between the

Dlog fO2 values obtained from BR219 (-1.27 log-bar

units) and those from GP40 and GP42, representative of

Greene Point unmetasomatized lithospheric mantle (-1.35

and -1.20 log-bar units, respectively).

Partial melting should lead to reduction of residual

mantle peridotite due to the preferential partitioning of

Fe3? into the produced melt (Bryndzia and Wood 1990;

Arculus 1994; Kadik 1997; Lee et al. 2003; McCammon

and Kopylova 2004 and reference therein); this in turn

should yield a negative correlation between fO2 and indi-

cators of the residual nature of peridotite (i.e., Cr# in spi-

nel). Contrary to this prediction, no clear correlations could

be found between fO2 values and the degree of depletion in

NVL peridotites. Rather, the samples from Greene Point

show a roughly mutual increase of Dlog fO2 and spinel Cr#.

Within this suite, the unmetasomatized GP42 and GP40

samples whose degree of partial melting F was estimated to

be 2 and 8 %, respectively, indicate virtually no change of

the local fO2 induced by partial melting event process

(Table 7). The other selected GP samples which have a

higher refractory character (harzburgites GP6, GP22 and

GP3), according to what stated above, should document a

local decrease of oxidation state; in contrast, they yield

values of Dlog fO2 higher than those obtained from GP42

and GP40 (Fig. 5). It should be noted that these harzburgite

samples are also those cryptically metasomatized, whose

spinel analysis provided the highest degree of oxidation

(z = 19–26 %; Fig. 5). This finding suggests that, although

no apparent correlation between estimated Dlog fO2 values

and the enrichment in incompatible elements could be

found, the high Dlog fO2 values obtained for GP metaso-

matized xenoliths could have been induced by the Ceno-

zoic metasomatism.

Several authors (Menzies et al. 1987; Ballhaus 1993 and

reference therein) have noted a fO2 versus Cr# correlation

in subcontinental mantle xenoliths, similar to that defined

by GP samples. The correlation has been ascribed to a

higher activity of metasomatizing oxidized fluids/melts on

a refractory harzburgite rather than on a fertile lherzolite

‘‘even if the amount of oxygen or Fe3? added to fertile and

refractory domains is the same’’ (Ballhaus 1993). Ionov

and Wood (1992), instead, did not detect any relationship

between fO2 values and the degree of depletion and/or

metasomatism in spinel peridotites from continental rift

environments. Nevertheless, within our samples, we

observe that peridotite xenoliths with comparable degree of

depletion (e.g., the unmetasomatized GP40 with F = 8 %,

the metasomatized GP31 or BR218* with F = 10 and

9 %), show variations in Dlog fO2 (-1.35 log-bar units,

-0.65 log-bar units and -1.52 log-bar units, respectively)

that cannot be explained by the low fO2-buffering associ-

ated to the depleted character of these xenoliths. This

observation suggests that, when associated with variable

chemical enrichments, Cenozoic metasomatism induced

local oxidation within GP lithospheric mantle, whereas at

Baker Rocks, the redox state of lithospheric mantle was not

altered by the metasomatic event. Hence, we infer that the

metasomatic melts acting in the two sites were different in

fO2 conditions.

Oxygen fugacity of metasomatizing melts

At Greene Point, the distinctive metasomatic overprint

on trace element patterns points to a silicate metasomatic

agent resembling the host lava of the xenoliths (SAX20,

Perinelli et al. 2006); however, textural features of the

GP xenoliths and the fast magma ascent rates exclude

that the metasomatic event and the entrapment of the

xenoliths are contemporary. Nevertheless, host lava

SAX20 can be used to bracket the oxygen fugacity

during metasomatism.

In porphyritic rocks, the estimated fO2 value may not

represent the redox conditions of host lava at its origin;

indeed, it is related to conditions prevailing during the

crystallization; in fact, fractionation of olivine and/or

clinopyroxene, whose Fe3?/Fetot ratio is lower than that of

melt, can promote relative oxidation (Ballhaus 1993).

However, the low content of crystals (\10 % vol. pheno-

crysts, Orlando et al. 1997) of SAX20 allows a reliable

estimate of the upper limit of fO2 using the Kress and

Carmichael (1991) model. The Fe2? and Fe3? content in

the melt has been derived from the experimental data on

Cr-spinel-silicate melt equilibria through the equation

proposed by Maurel and Maurel (1982):

log Fe2þ=Fe3þ� �sp¼ 0:764� log Fe2þ=Fe3þ� �

melt�0:343

ð16Þ

Using the compositions of chromium-bearing spinels

(Cr# = 0.63) in olivine phenocrysts (Fo = 78–81), the

calculated value of fO2 in SAX20, expressed as Dlog fO2,

ranges from ?1.1 to ?1.5 log-bar units, at T = 1095 �C

(olivine-spinel geothermometer, Ballhaus et al. 1991) and

P = 1.3 GPa (the mean value of equilibrium pressures

estimated for the GP xenoliths). The relatively high oxygen

fugacity inferred for this possible metasomatic melt(s) is


123

compatible with the increase of about 1 log-bar unit

observed in the metasomatized GP xenoliths (Fig. 6).

At Baker Rocks, the crystallization of amphiboles in the

metasomatic vein in the BR218 xenolith points to redox

conditions of Dlog fO2 = -1.71 log-bar units; this value

fairly matches the fO2 conditions estimated for the peri-

dotite matrix adjacent to the vein (BR218*; Fig. 6) and

suggests a similar oxidation state for both metasomatizing

melt(s) and the lithospheric mantle in this area.

Coltorti et al. (2004) proposed that a melt similar to

nephelinite SAX20 could represent a possible metasomatic

agent also at Baker Rocks. However, the experimental

study of Perinelli et al. (2008) did not provide any clear

evidence in support of this hypothesis because: (1) the

changes induced in the composition of primary minerals

only in part matched those found in minerals from natural

metasomatized xenoliths and (2) no run produced an

amphibole-bearing paragenesis.

Moreover, the high oxygen fugacity conditions esti-

mated for SAX20 metasomatic melt (Dlog fO2 = ?1.1 to

?1.5 log-bar units) are in contrast with those indicated by

both the amphibole-bearing vein in BR218 xenolith

(Dlog fO2 = -1.71 log-bar units) and the peridotite matrix

(Dlog fO2 = -1.52 log-bar units close to the vein, and

-1.14 log-bar units far from the vein). Therefore, at Baker

Rocks, a SAX20-like metasomatic silicate melt can be

excluded.

At Baker Rocks, the lack of Cr-spinel phenocrysts in the

less evolved magmas erupted in the Mt. Melbourne Vol-

canic Province prevents the application of the approach

used for SAX20 host lava to derive the fO2 conditions of

the lithospheric mantle. However, the titanomagnetite-

ilmenite pairs occurring in the groundmass of basanites/

olivine-basalts erupted in the area (Armienti et al. 1991)

can be used for this purpose. In fact, the application of the

titanomagnetite-ilmenite oxybarometer provides an esti-

mation of the range of fO2 at which basic magmas evolved

and hence provides an upper limit to the oxidation state of

shallow mantle beneath Baker Rocks. The Dlog fO2, cal-

culated by the titanomagnetite-ilmenite oxybarometer

proposed by Sauerzapf et al. (2008), varies from -1.6 to

-0.9 log-bar units, close to the range estimated for both

BR xenoliths and amphibole-bearing vein and far from

Dlog fO2 estimated for SAX20. These results strongly

support the conclusion that different melts are responsible

for the distinct styles of metasomatism of lithospheric

mantle of Greene Point and Baker Rocks recorded by the

xenoliths.

In summary, we can suggest that at Greene Point, a

SAX20-like melt metasomatized the peridotite portions

closest to the magma conduits (vein and/or dyke); in

contrast, at Baker Rocks, the metasomatism can be

described as a continuous differentiation process of ba-

sanitic/olivine-basalt melts that during their rise within

the lithospheric mantle, generated anhydrous and hydrous

veins plus Fe–Ti enrichment or cryptic enrichment of

incompatible elements in the surrounding peridotite. This

is in agreement with the mechanisms of formation of

hydrous/anhydrous metasomatic veins in the lithospheric

mantle proposed by several authors (i.e., Harte et al.

1993) and with the experimental studies of Pilet et al.

(2009), which suggest that the formation of amphibole-

bearing metasomatic veins can be explained by the

percolation and differentiation of small-melt fraction,

produced by low-degree partial melting of a depleted

mantle source.

Metasomatism and geodynamic context

As stated previously, the metasomatic event documented in

GP and BR mantle xenoliths is related to magmatic activity

during the Cenozoic Ross Sea rifting. The development of

the Ross Sea rift system and associated magmatic activity

in northern Victoria Land had been linked to dextral

transtensional movements connected with an important

reorganization in plate kinematics in the Southern Hemi-

sphere (Rocchi et al. 2002, 2005; Storti et al. 2007). Iso-

topic and geochemical features of NVL Cenozoic basalts

point to a model for which the source of these near-primary

melts is a sublithospheric mantle, metasomatized during

the initial extensional event that affected the West Ant-

arctic Rift System in the Late Cretaceous (Nardini et al.

2009). After the first phase of Ross Sea opening, at the

edge of the lithospheric step between the thick East Ant-

arctic craton and the thinned crust of the Ross Sea, an

‘‘edge effect’’ in mantle circulation (Faccenna et al. 2008)

may have provided the heat source to warm up the

-3.0

-2.5

-2.0

-1.5

-1.0

-0.5

0.0

0.5

1.0

1.5

GP6

GP Host lava

GP22

GP31

GP40

GP42

BR213

BR219

BR218

BR218*

BR218 Amph

Δlo

gfO

2

FMQ buffer

cryptically metasomatized

unmetasomatized

close to amphibole-bearing vein

crypticallymetasomatized

SAX20

Fig. 6 A summary of oxygen fugacity values estimated for Greene

Point and Baker Rocks mantle samples. fO2 values, expressed as

deviation from FMQ buffer, estimated from amphibole of BR218

xenolith and from the SAX20 melanephelinite are reported


123

lithospheric mantle beneath northern Victoria Land

(Armienti and Perinelli 2010).

During Eocene–Oligocene, the reactivation of Palaeo-

zoic lithospheric discontinuities in northern Victoria Land

was due to the increase in differential velocity across the

Southern Ocean fracture zones (Rocchi et al. 2002, 2005;

Storti et al. 2007). This induced local mantle decompres-

sion and melting providing magma that initially was em-

placed as plutons and dyke swarms (Rocchi et al. 2002).

From Late Miocene to Present, the craton-ward mantle

flow led to collapse of the rift shoulder and normal faulting

which favored the rise of magmas to the surface (Nardini

et al. 2009).

In this frame a ‘‘short-term metasomatic’’ mechanism

(Pilet et al. 2004, 2008) has been proposed to account for

the Sr–Nd–Pb isotopic ratios of NVL near-primary melts

(Nardini et al. 2009) and for isotopic heterogeneity of

parental and metasomatizing melts revealed by the geo-

chemistry of Browning Pass cumulates (Perinelli et al.

2011). The ‘‘short-term metasomatic’’ model envisages a

vein-plus-wall mechanism similar to that proposed by Fo-

ley et al. (2006) for magmatism linked to the Phanerozoic

Lambert-Amery Rift in eastern Antarctica, in which melts

produced during decompression mantle melting in the early

phase of rifting form veins in overlying lithosphere by a

percolative fractional crystallization process. These enri-

ched veins are probably more oxidized with respect to

surrounding depleted mantle. As the rifting proceeds, the

progress of mantle decompression and/or the onset of local

thermal perturbation promote melting of the low-solidus

enriched veins characterizing the chemistry of the early

magmas. While the rifting advances, more and more heat is

provided by the upward movement of isotherms and by

convection inducing the melting in the relatively reduced

peridotite matrix.

Thus, the differences in oxygen fugacity caused by

distinct metasomatic melts acting in Greene Point and

Baker Rocks, possibly reflect the contributions of two end-

members. The former represents the melting of metaso-

matic veins/domains emplaced during the amagmatic phase

of Ross Sea rifting, while the latter is the effect of partial

melting induced in the mantle wedge by edge-driven con-

vection of asthenospheric mantle, developed at the

boundary between thick East Antarctic Craton and thin

Ross Sea lithosphere.

Acknowledgments We wish to thank both the Editor, C. Ballhaus,

and the two reviewers, I. Parkinson and anonymous, for their careful

work and their very constructive comments. We gratefully acknowl-

edge funds provided by the project PRIN 2008 ‘‘SPIN GEO-TECH’’

and PRIN 2009 ‘‘Structure, microstructures and cation ordering: a

window on to geological processes and geomaterial properties’’.

M. Serracino (CNR-IGAG) kindly helped during EMP analyses.

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