Download Lithium Isotope Variations in Ocean Island

JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 11
PAGES 2333^2347
2012
doi:10.1093/petrology/egs052
Lithium Isotope Variations in Ocean Island
BasaltsçImplications for the Development of
Mantle Heterogeneity
M.-S. KRIENITZ1, C.-D. GARBE-SCHO«NBERG1, R. L. ROMER2*,
A. MEIXNER2y, K. M. HAASE3 AND N. A. STRONCIK2z
1
INSTITUT FU«R GEOWISSENSCHAFTEN, UNIVERSITA«T KIEL, OLSHAUSENSTR. 40, 24118 KIEL, GERMANY
2
DEUTSCHES GEOFORSCHUNGSZENTRUM, TELEGRAFENBERG, 14473 POTSDAM, GERMANY
3
GEOZENTRUM NORDBAYERN, SCHLOSSGARTEN 5, 91054 ERLANGEN, GERMANY
RECEIVED JULY 1, 2011; ACCEPTED JULY 10, 2012
ADVANCE ACCESS PUBLICATION AUGUST 21, 2012
Lithium elemental and isotopic compositions of 33 glass and
whole-rock samples from nine oceanic island regions were determined
to characterize the Li inventory of the deep mantle. The Li contents
of the investigated lavas range from 1·5 to 13·3 mg g1, whereas
d7Li ranges from 2·4 to 4·8ø. There are weak co-variations between the Li/Y, d7Li, and Sr^Nd^Pb isotope compositions of the
lavas, indicating that the Li elemental and isotopic characteristics
of ocean island basalt to some extent reflect mantle source heterogeneity. In detail, HIMU-type lavas are characterized by d7Li values
(up to 4·8ø) slightly heavier than those for average normal
mid-ocean ridge basalt (3·4 1·4ø) and by comparatively low Li
contents; EM1-type lavas are characterized by isotopically light Li
(average 3·2ø) and relative Li enrichment, whereas EM2-type
lavas tend to heavier d7Li values (up to 4·4ø) with high Li concentrations. The Li contents and isotope characteristics of
HIMU-type lavas are consistent with recycling of altered and dehydrated oceanic crust, whereas those of the EM1-type lavas can be
attributed to sediment recycling. The Li characteristics of
EM2-type lavas may reflect reworking of mantle wedge material
that has been infiltrated by fluids derived from the subducting plate.
KEY WORDS:
EM; HIMU; mantle reservoirs; MORB; OIB; lithium
I N T RO D U C T I O N
The chemistry and isotopic composition of oceanic basalts
provide fundamental information about the composition of
the Earth’s mantle (e.g. Alle'gre,1982; Zindler & Hart,1986;
*Corresponding author. Present address: Deutsches GeoForschungs
Zentrum GFZ, Telegrafenberg, 14473 Potsdam, Germany. Telephone:
þ49 (0) 331 288 1405. Fax: þ49 (0) 331 288 1474. E-mail: rolf.romer@
gfz-potsdam.de.
y
Present address: Universita«t Bremen, Fachbereich Geowissenschaften
(FB5), Bibliothekstr.1, 28358 Bremen, Germany.
Hofmann,1988,1997). Although the chemistry of the mantle
generally appears to be controlled by processes such as
magma generation and crust formation, mantle convection
and material recycling, as well as metasomatism, the relative
contribution of these processes to the heterogeneity of the
mantle is still debated. Combined studies of ocean island basalts (OIB) and mid-ocean ridge basalts (MORB) provide insights into dynamic processes within the Earth’s mantle,
because the chemistry and isotopic composition of OIB and
MORB closely reflect the compositions of their mantle
sources. Whereas MORB are generated in the upper, relatively trace element depleted mantle, OIB magma sources
probably reside in deeper mantle regions.
The compositional range of the mantle is defined by a
series of components that are thought to be the product of
different types of crust^mantle interaction during subduction (e.g. EM1, EM2, HIMU, DMM; Zindler & Hart,
1986; Hart, 1988; Hofmann, 1988; Sun & McDonough,
1989; Weaver, 1991). Subducted oceanic crust and pelagic
sediments, for example, or metasomatized subcontinental
lithosphere plus lower continental crust, may contribute to
the EM1 component, whereas EM2 probably contains
recycled oceanic crust and terrigenous sediments or
recycled portions of upper continental crust (e.g. Pilet
et al., 2005; Jackson et al. 2007; Workman et al., 2008). The
HIMU signature probably is the result of recycling of
altered and subduction-modified ancient oceanic crust
(e.g. Hofmann, 1997; Stracke et al., 2005), whereas the
z
Present address: Integrated Ocean Drilling Program, Texas A&M
University,1000 Discovery Drive, College Station,TX 77845-9547, USA.
The Author 2012. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
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JOURNAL OF PETROLOGY
VOLUME 53
DMM component is characterized by incompatible trace
element depletion that is thought to represent a complementary reservoir to the trace element enriched continental crust (e.g. Hofmann, 1988; Sun & McDonough, 1989).
As lithium contents and isotopic compositions are variable in continental crust, oceanic crust, sediments, and
the Earth’s mantle, lithium can be used as a tracer for material transport from the surface into the deep mantle. In
this regard, lithium and particularly its isotopes have attracted interest in recent years to investigate subduction
zone and recycling processes (e.g. Tomascak et al., 2000;
Benton et al., 2004; Elliott et al., 2004, 2006; Halama et al.,
2008; Vlaste¤lic et al., 2009). The trace element lithium is a
light alkali metal, moderately incompatible during mantle
melting, and thus concentrated in the crust relative to the
mantle. Lithium is mobile in aqueous fluids and has two
stable isotopes, 6Li and 7Li, that are, unlike the radiogenic
Sr^Nd^Pb isotope systems, unaffected by time-dependent
parent^daughter fractionation processes. The preferential
loss of 7Li during weathering results in isotopically light
crust and isotopically heavy river water and seawater (e.g.
Chan et al., 1992, 2002; Elliott et al., 2004). In marine environments Li isotopes fractionate during low-temperature alteration of oceanic crust, which will be enriched in 7Li
relative to fresh rocks. The isotopic composition of Li
therefore is a sensitive tracer for seawater contamination
(e.g. Chan et al., 1992, 2002; Decitre et al., 2002; Tomascak
et al., 2008; Scholz et al., 2009, 2010). During subduction,
oceanic crust and sediments are recycled into the mantle.
Dewatering and devolatilization processes transfer Li
from the slab to the mantle wedge. Metamorphic dehydration of the slab and related isotopic fractionation produce
residual slab rocks that are depleted in 7Li (Chan et al.,
1993, 2002; Seyfried et al., 1998). The progressive
fluid-related loss of Li during subduction makes slab Li
progressively lighter with deeper subduction (e.g. Zack
et al., 2003). Fluids released from the slabçalthough carrying heavier Li than the slabçwill be increasingly lighter
with increasing distance from the subduction zone (e.g.
Agostini et al., 2008). Thus, the transfer of slab Li to the
mantle wedge may be spatially variable both in amount
and isotopic composition (Seyfried et al., 1998; Chan et al.,
2002). Because isotopic fractionation of Li appears to be
small at magmatic temperatures, Li isotope ratios are not
significantly affected by melting processes or magmatic
differentiation (Tomascak et al., 1999b; Chan & Frey, 2003;
Halama et al., 2007). Earlier studies on MORB demonstrated a relatively homogeneous Li isotopic composition
of the depleted mantle reservoir with d7Li of þ1·5ø to
þ5·6ø (e.g. Tomascak, 2008; Tomascak et al., 2008, and
references therein). Carbonatites, sampling the deep
mantle, have been shown to sample a similarly homogeneous mantle reservoir with d7Li of þ3·3ø to þ5·1ø
(e.g. Halama et al., 2007) that seems not to have changed
NUMBER 11
NOVEMBER 2012
its Li isotopic composition through time (e.g. Halama
et al., 2008). This coincidence of d7Li in MORB and carbonatites may indicate that upper mantle rocks encompass
a relatively small compositional range. Evidence for Li isotopic heterogeneity in the mantle mainly comes from
mantle xenoliths and OIB. Xenoliths from mantle that
was metasomatized by fluids or melts released from
deeply subducted rocks may show distinctly negative d7Li
values (e.g. Nishio et al., 2004) and their combined Li^Sr^
Nd isotope systematics indicate that the EM1 reservoir
might have a rather light Li isotopic composition (see
Nishio et al., 2004). Although the range of d7Li values of
OIB (þ2ø to þ8ø, e.g. Chan & Frey, 2003; Tomascak,
2008; Vlaste¤lic et al., 2009) significantly overlaps with the
MORB range, the high d7Li values in some OIB suggest
that their formation involved a mantle reservoir with distinctly heavier Li isotopic compositions than MORB.
In this contribution we report Li contents and isotopic
compositions for volcanic glasses and whole-rocks from different oceanic island regions. The main objectives are to
examine the Li isotope characteristics of different mantle
reservoirs, as defined by the isotope systems of Sr, Nd,
and Pb, and to infer the role of material recycling in the
development of mantle heterogeneity.
SA M P L E S ET
The sample set consists predominantly of submarine basalts from nine hotspot settings, including the Azores,
Macdonald, Pitcairn, Re¤union, St. Helena, Society islands
and the Juan Fernandez, Easter and Foundation Seamount
Chains (Fig. 1a). The magmatism in all of these regions is
thought to be generated by mantle plumes that sample
mantle domains with different geochemical and isotopic
compositions (Fig. 1b and c). The sampled regions were selected to encompass a large part of the Sr^Nd^Pb isotopic
compositional range of OIB and to be as close as possible
to the various isotopic mantle components of Zindler &
Hart (1986). The EM1 mantle source component is represented by lavas from Pitcairn (Woodhead & Devey, 1993;
Eisele et al., 2002), whereas lavas from the Society (Devey
et al., 1990) and Re¤union hotspots (Fretzdorff & Haase,
2002) have characteristics of an EM2 mantle component.
A HIMU mantle source signature is characteristic for
lavas from the St. Helena region (Chaffey et al., 1989). The
Macdonald seamount is the youngest expression of the
Austral plume track and probably contains small amounts
of sediment in its source region (Hemond et al., 1994).
Lavas from the Easter and Foundation Seamount Chains
represent mixtures between DMM and a HIMU component (Hemond & Devey, 1996; Cheng et al., 1999), whereas
lavas from Juan Fernandez isotopically lie between HIMU
and prevalent mantle (Zindler & Hart, 1986) compositions
(Devey et al., 2000). Compositions of subaerial lavas from
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KRIENITZ et al.
LITHIUM ISOTOPES IN OIBs
(a)
Azores
ty
e
ci
So
na
o
cd
n
a
M
F
0.5133
Ea
St. Helena
z
na
n
r
Fe
a
Ju
0.708
DMM
(b)
0.5131
Pitcairn
Society
Réunion
Azores
Macdonald
St Helena
Easter SC
Foundation SC
J Fernandez
(c)
EM2
0.707
0.706
0.5129
87Sr/86Sr
143Nd/144Nd
Réunion
e
nd
SC
io
at
d
n
ou
S
er
st
t
Pi
ld
C
rn
i
ca
EM1
HIMU
0.705
0.5127
0.704
EM2
0.5125
HIMU
EM1
0.5123
0.702
0.703
0.704
0.705
DMM
0.706
0.707
0.708 16
87Sr/86Sr
17
18
19
20
21
0.703
0.702
22
206Pb/204Pb
Fig. 1. (a) General map of the nine hotspot localities from which OIB lavas were analyzed. Easter SC, Easter Seamount Chain; Foundation SC,
Foundation Seamount Chain. (b, c) Sr^Nd^Pb isotope compositions of the OIB lavas. Data represent whole-rock analyses except for the
Easter Seamount Chain and Re¤union data, which are glass analyses. Mantle reservoir compositions (DMM, EM1, EM2, HIMU) are from
Hart (1984), Zindler & Hart (1986), Hart et al. (1992) and Stracke et al. (2005). (See Electronic Appendix Table EA1 for data sources.)
Sa‹o Miguel, Azores, are explained by a mixture between
DMM and enriched mantle sources probably containing
recycled oceanic crust and isotopically falling between
EM2 and HIMU compositions (Beier et al., 2007; Elliott
et al., 2007; Fig. 1b and c).
M ET HODS
Fresh glasses were separated by careful handpicking using
a binocular microscope. After picking the glass shards
were cleaned in purified ethanol for several minutes and
used for subsequent analyses. Whole-rock compositions
were determined only for four lavas from the Azores
(Beier et al., 2007) and those samples were cut into pieces,
crushed, sieved, ultrasonically cleaned with purified water
and thereafter powdered in an agate ball mill.
A subset of samples was analyzed for their major and
trace element content. Major element compositions of
glasses were determined by standard wavelengthdispersive electron microprobe analysis at Deutsches
GeoForschungsZentrum GFZ, Potsdam (Germany) using
a CAMECA SX-100 instrument and at Institut fu«r
Geowissenschaften, Universita«t Kiel (Germany) with a
JEOL Superprobe 8900 electron microprobe. A set of
international mineral reference materials was employed
for calibration and the instruments were operated at an
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JOURNAL OF PETROLOGY
VOLUME 53
accelerating voltage of 15 kV and a 15 nA beam current
with a defocused beam (e.g. Trumbull et al., 2003; Stroncik
et al., 2009).
Trace element analyses were performed at the Institut
fu«r Geowissenschaften, Universita«t Kiel (Germany) with
an upgraded VG PlasmaQuad PQ1 inductively coupled
plasma (ICP)-mass spectrometer following the analytical
procedure of Garbe-Scho«nberg (1993). Comparison of duplicate digestions of the same sample gave a standard deviation typically better than 1%. The accuracy of the data
based on reference material BHVO-2 is better than 6%
for most of the elements and can be estimated from analyses of BHVO-2 shown in Electronic Appendix Table
EA1 together with the results of the major and trace element determinations as well as published geochemical and
isotope data for the sample set (the Electronic Appendices
are available for downloading at http://www.petrology
.oxfordjournals.org).
Lithium abundances and isotopic compositions of the
samples were determined at Deutsches GeoForschungs
Zentrum GFZ, Potsdam (Germany). Between 15 and
50 mg of sample material were digested in a mixture of
concentrated hydrofluoric and nitric acid, dried, and
re-dissolved in 6 ml1N nitric acid and 3 ml methanol.
Reference material and procedure blank were run with
each sample series.
Ion exchange techniques used to produce, isolate and
purify Li principally follow the methods of Tomascak
et al. (1999a) and Jeffcoate et al. (2004).
Li was separated using quartz columns with 10 ml
AG-50 X8 cation exchange resin and 1N HNO3 in
80% methanol as eluent. The sample was collected in
103 ml eluate, an aliquot of which was used for Li concentration determination. Li eluates with significant
Na contents were processed through a second cleaning
step using 0·5 N/1N HCl:80% methanol (Jeffcoate
et al., 2004). Possible loss of lithium during the separation procedures, which could cause isotopic fractionation (Taylor & Urey, 1938; Tomascak et al., 1999a),
was monitored by scanning the Li concentration of
12 ml eluate before and 5 ml (0·5 ml) eluate after the
main Li fraction (second column step: 2 ml pre-Li eluate
and 0·5 ml post-Li eluate fractions). The total loss of lithium was always less than 0·1% of the total amount of Li
in the samples and thus had a negligible effect on the isotopic composition of the sample. The influence of the procedural blanks (between 10 and 280 pg Li) was
insignificant and did not exceed 0·1% of the sample Li.
Li isotopic compositions and contents were determined on
a ThermoFisher Scientific NEPTUNE multicollector
(MC)-ICP-mass spectrometer. All samples were measured
repeatedly in the standard/sample/standard bracketing
mode. Data collectionand reduction followedthe procedures
described by Wunder et al. (2006, 2007). Samples and
NUMBER 11
NOVEMBER 2012
standards were diluted in 2% HNO3 and closely adjusted to
25 ppb Li (5%) to avoid bias from variable concentrations
onthebracketing. Before andaftereach sample and standard,
2% HNO3 was measured, andthe average of these two measurements was used as an instrumental baseline to correct the
respective sample or standard. Lithium isotopic compositions are reported as d7Li (d7Li ¼ {[(7Li/6Li)sample/
(7Li/6Li)reference material] ^ 1} 1000) relative to the reference
material NIST SRM 8545 (L-SVEC). Li elemental and
isotopic data for international reference materials analyzed
during this study are presented in Electronic Appendix
Table EA2. Repeated isotope analyses of thevariousreference
materials give an analytical precision of 0·2ø at the 2s
level and our results of the standard analyses fall within the
ranges of previously published values (Electronic Appendix
Table EA2).The Li content and Li isotope composition of the
samples are reported inTable1.
R E S U LT S
Lithium concentrations in the OIB lavas range from 1·5 to
13·3 mg g1 and typically increase with decreasing MgO
content as expected on the basis of the moderate incompatibility of Li (Fig. 2a). In general, OIB have lower Li
concentrations for a given MgO than average MORB
(Chan et al., 1992; Tomascak et al., 2008). Most of the lavas
from St. Helena, Easter Seamount Chain, Foundation
Seamount Chain and Juan Fernandez have Li/Y ratios
50·20, whereas lavas from Pitcairn, Society, Re¤union,
Azores and Macdonald have typically higher Li/Y ratios
than the 0·20 characteristic of MORB (Fig. 2b). One
sample of the Macdonald group (64 DS-2) yields a significantly lower Li concentration of 1·5 mg g1 and thus a
much lower Li/Y ratio (0·06) than other lavas with comparable MgO contents (Fig. 2b). There is no petrographic
evidence for alteration (e.g. alteration rims, palagonitization, smectite) and no geochemical evidence for alteration
(e.g. unusual Na or K contents compared with the rest of
the MacDonald samples). Furthermore, the Li isotopic
composition of this sample overlaps, within error, the
other samples of Macdonald OIB lavas (Figs 3 and 4).
Variations of Li/Y versus Ba/Rb, Cl/K or Sr^Nd isotopic
composition are not coherent for all OIB groups (Fig. 2).
For example, negative correlations between Li/Yand Cl/K
are observed for Society group lavas, whereas lavas from
the Macdonald seamount are slightly positively correlated,
and the remaining groups show no obvious co-variation
(Fig. 2d).
The results of the lithium isotope determination are
shown in Fig. 3. The d7Li values range from 2·4 to 4·8ø
and overlap entirely with the average MORB value of
3·4 1·4ø (Tomascak et al., 2008). The data are comparable with published values from other OIB worldwide (e.g.
Chan & Frey, 2003; Ryan & Kyle, 2004; Chan et al., 2006a),
although our data fall in a much narrower range than the
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KRIENITZ et al.
LITHIUM ISOTOPES IN OIBs
Table 1: Lithium concentrations and Li isotopic compositions of the OIB samples analysed in this study
Sample
Cruise
Location
Li (mg g1)
d7Li (ø)
46 DS-2
SO 65
Pitcairn
6·6
3·3
51 DS-1
SO 65
Pitcairn
5·7
3·2
66 DS-3
SO 65
Pitcairn
5·4
3·4
9 DS-2
SO 47
Society
6·5
3·9
122 DS-2
SO 65
Society
12·4
3·3
122 DS-4
SO 65
Society
12·4
3·2
22 GTVC-1
SO 47
Society
7·8
4·1
17 DS-2
SO 87
Réunion
6·8
3·9
18 DS-2
SO 87
Réunion
5·9
4·4
59 DS-3
SO 84
St. Helena
13·3
4·8
59 DS-4
SO 84
St. Helena
12·9
4·0
63 DS-1
SO 84
St. Helena
6·5
3·6
63 DS-3
SO 84
St. Helena
8·0
4·7
60 GTVA-2
SO 47
Macdonald
7·9
3·9
64 DS-1
SO 47
Macdonald
5·3
3·8
64 DS-2
SO 47
Macdonald
1·5
4·2
110 DS-1
SO 65
Macdonald
5·6
3·9
110 DS-6
SO 65
Macdonald
5·6
3·5
11 DS-1
SO 100
Foundation SC
5·7
3·2
70 DS-2
SO 100
Foundation SC
6·0
3·3
97 DS-2
SO 100
Foundation SC
3·9
2·4
99 DS-1
SO 100
Foundation SC
4·6
3·3
25 DS-3
SO 80
Easter SC
6·3
4·0
37 DS-1
SO 80
Easter SC
5·9
4·1
39 DS-2
SO 80
Easter SC
4·9
3·6
27 DS-2
SO 80
Easter SC
4·9
3·4
43 DS-1
SO 80
Easter SC
4·9
2·6
SM 0101
—
Azores
4·1
4·5
SM 0134
—
Azores
10·8
3·2
SM 9704
—
Azores
4·9
4·2
SM 9716
—
Azores
9·5
3·3
7 DS-1
SO 80
Juan Fernandez
7·7
2·9
10 DS-5
SO 80
Juan Fernandez
8·7
3·5
same d7Li values at variable Sr and Nd isotopic composition. Similarly, lavas from the Foundation Seamount
Chain have variable Sr^Nd^Pb isotopic compositions, but
little variation in their d7Li (Fig. 5). In contrast, lavas
from St. Helena, MacDonald, Juan Fernandez and
Re¤union show only small variations in their Sr^Nd^Pb
isotopic composition, whereas their d7Li value varies significantly (Fig. 5).
DISCUSSION
Secondary alteration and lithium isotope
signals
SC, seamount chain.
published data (see Figs 4 and 5). Taking the Li isotopic
composition of average normal (N)-MORB as reference,
lavas from St. Helena and Re¤union tend to higher values
and those of the Pitcairn and the Foundation Seamount
Chain groups are all slightly lower. Li isotopic compositions of the remaining OIB suites scatter in the field of
MORB (Fig. 3). Although there seem to be systematic, but
not necessarily coherent, variations of d7Li versus MgO,
Li, Li/Y or Sr^Nd^Pb isotopes for the various locations,
there is no consistent pattern for all OIB as a group (Figs
4 and 5). For instance, Pitcairn group samples have the
A prerequisite for the investigation of the primary magmatic signatures of lavas is the identification and omission
of samples that have been modified by secondary processes. This is especially crucial for lithium and its isotopes
as the variation of d7Li is likely to be relatively small and
the d7Li of seawater (d7Li 31ø; e.g. Millot et al., 2004)
is much higher than that of typical oceanic lavas.
Interaction with seawater for submarine lavas or weathering for subaerial samples may effectively modify the primary Li contents and isotopic signatures. It is well
established that, at low temperatures, reactions between
seawater and basalt result in Li enrichment and heavy
d7Li signatures in the rocks (e.g. Chan et al., 1992, 2002;
Moriguti & Nakamura, 1998). Figure 6 shows the effects of
potential seawater alteration on the OIB lavas. Owing to
the large compositional difference between the reservoirs,
mixing trends between the components are strongly
curved. Thus, for a small addition of seawater Li, the d7Li
values in the rocks will increase significantly, whereas the
Sr^Nd^Pb isotopic composition shows little or no variation. For instance, St. Helena group lavas form vertical
trends when d7Li is plotted versus Sr^Nd^Pb isotopic composition (Figs 5 and 6). Although the various OIB group
trends seem to coincide with the variation expected for
post-magmatic rock^seawater interaction, these trends are
not sufficient to prove seawater alteration, as subduction
of altered oceanic crust would produce the same pattern
in the OIB source. The best evidence for post-magmatic alteration by seawater may be obtained from the ratio Cl/
K. Recent interaction with seawater would produce high
Cl/K in the rocks at relatively constant d7Li. Such a pattern is not observed (Fig. 6c). Based on these arguments
and because most of our samples are fresh, carefully
hand-picked glasses we consider the observed variability
in Li content and isotopic composition of the OIB lavas as
not being produced by secondary alteration processes. It
should be noted that, as not all literature data have been
checked as thoroughly for post-eruption alteration as the
data presented in Table 1, the possibility of secondary processes accounting for the larger d7Li range in the literature
data cannot be excluded.
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Fig. 2. MgO vs (a) Li concentration and (b) Li/Y, and Li/Y vs (c) Ba/Rb, (d) Cl/K, (e) 87Sr/86Sr, and (e) 143Nd/144Nd of basalts from the various OIB groups. MORB values (filled star) in (a)^(c) are from Nishio et al. (1999, 2007) and Tomascak et al. (2008). The average East Pacific
Rise (EPR) MORB data (open star) in (c) is from Regelous et al. (1999), Elliott et al. (2006) and Tomascak et al. (2008).
Lithium isotope variation of OIB mantle
sources
Lithium isotope heterogeneity in the mantle
Our data demonstrate that the Earth’s mantle is heterogeneous in terms of its Li isotopic composition, although the
observed Li isotope variability in the studied OIB is small
and overlaps with the composition of average MORB
(3·4 1·4ø; Tomascak et al., 2008). Furthermore, combining the Li isotope data with Sr^Nd^Pb isotopes, which
have been used to characterize mantle components
(Zindler & Hart, 1986), it seems that different mantle reservoirs have different Li contents and Li isotopic signatures.
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KRIENITZ et al.
LITHIUM ISOTOPES IN OIBs
Fig. 3. Abundance histograms for the lithium isotope compositions of
OIB. The vertical line and the bordering shaded range represents the
average MORB d7Li value of 3·4 1·4ø (Tomascak et al., 2008). SC,
Seamount Chain. Error bar represents the typical 2s uncertainty of
our data.
Fig. 4. d7Li vs (a) MgO, (b) Li concentrations, and (c) Li/Y ratio of
OIB. Large symbols, data from Table 1; small symbols, additional literature data. Sources for additional data in (b): St. Helena: Ryan &
Kyle (2004); Raivaevae, Mangai, Rapa: Chan et al. (2009); Tahiti,
Marquesas, Samoa: Nishio et al. (2005); Re¤union: Albare'de &
Tamagnan (1988), Ryan & Kyle (2004); Pitcairn: Elliott et al. (2006);
MORB: Chan et al. (1992), Moriguti & Nakamura (1998), Elliott et al.
(2006), Nishio et al. (2007), Tomascak et al. (2008); Hawaii: Frey et al.
(1994), Roden et al. (1994), Norman & Garcia (1999), Chan & Frey
(2003). Error bars represent typical 2s uncertainty of data from
Table 1; some of the literature data have distinctly larger analytical
uncertainties.
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Although there are no clear trends suggesting mixing
between different mantle components, the lavas representing extreme compositions in Sr^Nd^Pb isotope space
show distinct Li isotope compositions. For example, the
St. Helena glass samples have the highest d7Li, which is in
agreement with previous findings on whole-rocks from
HIMU sources (Fig. 5). The composition of EM-type lavas
has not been defined well previously and the EM-type
glasses have d7Li between 3 and 4ø in agreement with
two previously analysed samples (Fig. 5). Trends representing mixing between DMM and HIMU source components
are displayed by lavas of the Easter Seamount Chain
group, whose d7Li values vary with radiogenic isotopic
composition (Fig. 5). Although the total variation in d7Li
is small, these trends are consistent with regard to different
mantle components and, as discussed in the previous section, cannot be explained by seawater alteration (Fig. 6).
Although it has been suggested that the Li isotope compositions of melts may be modified as a result of diffusive reactions between melt and wall-rock, the magnitude of this
effect is not well known and is likely to be small in systems
with high melt/rock ratios (Lundstrom et al., 2005;
Jeffcoate et al., 2007). Diffusive effects therefore are unlikely
to explain the observed Li isotope variability. Instead, the
Li^Sr^Nd^Pb isotopic correlations indicate that the d7Li
variability must be ascribed to mantle source
heterogeneity.
The lithium isotopic compositions of the three EM1-type
samples from Pitcairn are relatively uniform, averaging at
þ3·29 0·07ø (Fig. 5); that is, they are close to the average N-MORB value (Fig. 5). Based on the Li^Sr^Nd relations of the Pitcairn lavas (Fig. 5) and using the Sr^Nd^
Pb isotope signatures of the EM1 mantle source suggested
by Zindler & Hart (1986), an average d7Li value of about
þ2·7ø can be extrapolated for the EM1 mantle source,
which agrees well with published data from Pitcairn and
other EM1-type volcanic rocks (James & Palmer, 2000;
Chan & Frey, 2003; Ryan & Kyle, 2004). EM2-type lavas
from the Azores, Re¤union and Society have heavier Li isotope compositions than the EM1-type basalts, with the
highest observed d7Li value being about 4·4ø (Fig. 5).
Extrapolating the lithium isotopic range for these lavas to
the EM2 component Sr^Nd^Pb isotope composition suggested by Zindler & Hart (1986) gives d7Li values of
about 4·5^6·3ø for the EM2 mantle reservoir. With
regard to HIMU mantle sources, our results corroborate
studies of HIMU-related lavas (e.g. Jeffcoate & Elliott,
2003; Ryan & Kyle, 2004; Nishio et al., 2005) indicating
that a heavy d7Li signature (44·8ø) seems to be characteristic for this mantle component (Fig. 5).
Recycling of Li into the mantle
The subduction process recycles Li from surface reservoirs
(e.g. oceanic crust, sediments) into the Earth’s mantle.
During low-temperature alteration the oceanic crust will
NUMBER 11
NOVEMBER 2012
Fig. 5. (a^c) d7Li values of OIB vs Sr^Nd^Pb isotope composition
complemented by the isotopic compositions of whole-rocks and olivine separates from OIB and MORB. The average MORB value
(filled star) is taken from Tomascak et al. (2008), and the Sr^Nd^Pb
isotope composition of DMM from Zindler & Hart (1986). The d7Li
value of the various mantle end-members was defined by using the
same approach as Zindler & Hart (1986) when they defined the various end-members that are necessary to explain the Sr^Nd^Pb isotopic
variability of mantle-derived magmatic rocks. Thus, the radiogenic
isotopes were used as abcissa and the d7Li values were estimated in
such a way that (1) the variation between d7Li and the Sr, Nd, and
Pb isotopic composition is internally consistent and (2) the variation
of the Li isotopic composition between the various mantle
end-members remained small. Data from Table 1 and Electronic
Appendix Table EA1. (See caption of Fig. 4 for additional data
sources.) Error bars represent the typical 2s uncertainty of the data
in Table 1.
2340
KRIENITZ et al.
LITHIUM ISOTOPES IN OIBs
Fig. 6. d7Li values of OIB vs (a) 87Sr/86Sr, (b) 206Pb/204Pb, and (c)
Cl/K. Insets show the data on a larger scale. Mixing trends between
seawater and different rock compositions are indicated by black lines
and the numbers give the percentage of Li (a, b) and Cl (c) derived
from seawater. Seawater data (concentration and isotopic composition) are taken from Broecker & Peng (1982), Burke et al. (1982), Li
(1982, 1991), Chan & Edmond (1988), Ling et al. (1997) and Millot
et al. (2004). Error bars represent the typical 2s uncertainty of the
data in Table 1.
7
be enriched in Li owing to the uptake of heavy Li from
seawater (Chan et al., 1992, 2002). Because the lithium
transfer between fluids and rocks is temperaturecontrolled, alteration phases and minerals will be depleted
in Li at temperatures above 2008C, whereas associated
fluid phases are enriched in isotopically heavier 7Li
(Berger et al., 1988; Chan et al., 1992; Seyfried et al., 1998;
Pistiner & Henderson, 2003; Scholz et al., 2009, 2010). As
oceanic crust is transported into the deeper mantle it dehydrates as a consequence of compaction and temperature increase (Kerrick & Connolly, 2001). Lithium is transferred
from solid phases into fluids and expelled. Significant Li
isotope fractionation associated with this redistribution of
Li decreases with increasing temperature (Chan et al.,
1993, 2002). Currently there is much debate about the lithium isotopic composition of recycled oceanic crust. For instance, 7Li-enriched pore fluids expelled from the
accretionary prism of the Costa Rica subduction zone
(Chan & Kastner, 2000; Benton et al., 2004) and eclogites
with very low d7Li values (Zack et al., 2003) both indicate
that recycled oceanic crust probably is characterized by
very low d7Li values. Other studies, however, show that
eclogites may have highly variable d7Li and that large portions of the subducted Li can be retained in high-pressure
metamorphosed slabs (Marschall et al., 2007). In this case,
recycling of residual slab rocks could also contribute a
heavy d7Li signal to OIB mantle source regions. This
being said, it is likely that there is no single d7Li value for
the subducting oceanic slab, as the loss of heavy Li from
the slab depends on the dehydration history during subduction and the stability of Li-sequestering primary and
metamorphic mineral phases.
Fluids released from altered oceanic crust are likely to
have high Li concentrations and to carry a heavy d7Li
signal [e.g. modelled by Marschall et al. (2007)]. Mantle
wedge regions situated directly above descending plates
could therefore be infiltrated by such 7Li-enriched fluids
(Jeffcoate & Elliott, 2003). In this context, it has been
shown that arc mantle xenoliths have d7Li values similar
to or slightly heavier than MORB (Ionov & Seitz, 2008).
These xenoliths, however, represent fragments of the shallow mantle above a subduction zone and, therefore, do
not necessarily reflect the Li elemental and isotopic signature, which may be significantly different from it (Ionov
& Seitz, 2008). Budget calculations of arc regimes, however, indicate that the incorporation of slab-derived
mobile components into the source regions of arc basalts
is not quantitative (e.g. McCulloch & Gamble, 1991). This
indicates that a significant portion of these components
may be retained in the hydrated regions of the mantle
wedge (Ryan et al., 1995). Recycling of hydrated wedge material would therefore transport a heavy d7Li signal into
deeper parts of the mantle.
d7Li in the HIMU mantle component
The HIMU mantle component as represented by lavas
from St. Helena is characterized by high d7Li (Fig. 5). A
heavy d7Li mantle signature can be achieved either by
the addition of mantle wedge material to the OIB source
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JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 11
NOVEMBER 2012
region or by recycling of partially dehydrated altered oceanic crust. Although the recycling of mantle wedge material
gives a reasonable explanation for the development of
heavy Li isotopic mantle domains, it fails to explain the
characteristic incompatible element and Pb isotope compositions of HIMU lavas. The high Nb/La and Ce/Pb of
HIMU lavas, for example, indicate that their mantle
sources are depleted and not enriched in fluid-mobile elements, arguing against mantle wedge material as a major
reservoir contributing to HIMU sources (Fig. 7). Instead,
the elevated Li isotopic compositions in HIMU-related
lavas along with their relative depletion in fluid-mobile incompatible elements favour the hypothesis that HIMU
mantle sources contain recycled dehydrated oceanic crust
(Hofmann, 1997; Stracke et al., 2005).
d7Li and the EM1 signature in Pitcairn lavas
The EM1-type OIB show similarities in some incompatible
trace element ratios (e.g. La/Sm, Sr/Nd, Nd/Hf) to
HIMU OIB, indicating that both types of basalt might
share a common source component. This common mantle
component could be either subduction-modified, dehydrated oceanic crust or metasomatically overprinted
mantle rocks. Additionally, the EM domains may contain
sediments (Weaver, 1991; Woodhead & Devey, 1993;
Hemond et al., 1994; Plank & Langmuir, 1998; Eisele et al.,
2002), metasomatized mantle wedge material (Tatsumi,
2000; Lassiter et al., 2003; Niu & O’Hara, 2003), or delaminated continental crust (Willbold & Stracke, 2006).
Analyses of marine sediments reveal that they are the most
Li-enriched reservoir with a d7Li range for most samples of
0ø to þ7ø, but with a low global average of þ3ø and average concentration of 43·3 mg g1 Li (Chan et al., 2006b). For
lower continental crustal sections a concentration weighted
average d7Li value of 2·5ø with an average Li concentration
of 8 mg g1 is estimated (Teng etal.,2008).Thus, the addition
of either sediments or lower crustal material to altered oceanic crust with a calculated average d7Li value of 10ø
(Chan et al., 2002) would lower the overall Li isotopic composition. Subduction additionally modifies d7Li values, producing maximal 5ø lighter Li isotopic compositions
(Marschall et al., 2007). Li contribution from sediments or
lower crustal rocks therefore could account for the light
d7Li values observed in the EM1 basalts. Recycling of lower
crustal rocks, however, is unable to explain the relatively
high Li contents of EM1-type basalts because lower crustal
Li concentrations resemble that of altered oceanic crust
with a weighted average Li concentration of 7·6 mg g1
(Chan et al., 2002). Consequently, sediment recycling may
better account for the comparatively higher Li concentration at a given MgO content and low Li isotopic compositions of EM1-type basalts (Figs 2 and 5).
Fig. 7. d7Li values of OIB vs (a) Ce/Pb and (b) Nb/La. Error bars
represent the typical 2s uncertainty of the data from Table 1.
d7Li and the EM2 signature
The samples from the Society hotspot and one sample
from Sa‹o Miguel, Azores, with a high Sr and 206Pb/204Pb
isotope ratio represent the EM2 mantle component
(Fig. 1b). It has been suggested that upper continental
crust, terrigenous sediments, or mantle wedge material
contribute to EM2 mantle sources (e.g. Weaver, 1991;
Willbold & Stracke, 2006).
Although the relatively high Li contents of 35 mg g1 of
the upper continental crust (Teng et al., 2004) provide a
reasonable explanation for the Li enrichment in EM2 basalts, its low average Li isotope composition (2ø to
þ2ø; Teng et al., 2004) rules out its contribution to the development of EM2 mantle domains, because characteristically high d7Li values cannot be achieved after
subduction. Terrigenous sediments are also unlikely to contribute to EM2 mantle sources, as terrigenous sediments
then must have heavier d7Li values than pelagic material
(EM1), which is not the case. Analyses of marine sediments
indicate that pelagic material has variable d7Li values,
but typically has a heavier Li isotope composition
(1·3^14·5ø) compared with sediments derived from
2342
KRIENITZ et al.
LITHIUM ISOTOPES IN OIBs
terrestrial sources (1·7ø to 2·5ø; Bouman et al., 2004;
Chan et al., 2006b). This observation is inconsistent with
the recycling of different sediment types to explain the
chemical and isotopic differences of the EM mantle
source domains.
Alternatively, EM2 mantle signatures could be produced by mantle wedge recycling. This hypothesis is supported by the characteristic enrichment of fluid-mobile
and highly incompatible trace elements (e.g. Cs, Rb, Ba,
U, K) in EM2 OIB. This is also supported by the unradiogenic Os isotope compositions observed in EM2 lavas
from the Cook^Austral region (Lassiter et al., 2003),
whose youngest expression is the Macdonald seamount.
Unradiogenic Os isotope compositions are inconsistent
with sediment recycling in addition to altered oceanic
crust, but have been suggested to reflect mantle wedge
recycling (Lassiter et al., 2003). Elevated d7Li values (4·5^
5·6ø) of whole-rocks and olivine separates from basalts
from the same region and from Samoa (Chan et al.,
2009) with a strong EM2 signature strengthen this hypothesis. This model was also suggested by Chan et al.
(2009) (Fig. 5).
d7Li isotope signature of mantle end-members and constraints on Li input into the mantle
The variability of the Sr, Nd, and Pb isotopic composition
of OIB, which has been used to define geochemically distinct mantle reservoirs (Zindler & Hart, 1986), stands in
strong contrast to the small range of Li contents and Li isotopic compositions of mantle-derived rocks (see Fig. 5).
None the less, there are small variations in the contents
and isotopic composition of Li that show systematic relations with the isotopic compositions of Sr, Nd, and Pb,
indicating that Li in the mantle is not homogeneously distributed and there are subtle variations among the various
mantle reservoirs. These small variations among
mantle-derived rocks, furthermore, are in marked contrast
to the large range in both Li content and Li isotopic composition of the rocks being subducted and in part recycled
into the mantle. For instance, basaltic eclogites and metasomatized mantle xenoliths have d7Li values as low as 11
and 17, respectively (e.g. Zack et al., 2003; Nishio et al.,
2004). Actually, pyroxenite veins in peridotites ranging in
d7Li from þ11·8ø to 4·2ø (Brooker et al., 2004) indicate
that, on a small scale, the mantle locally is isotopically heterogeneous. It should be noted that mantle-derived rocks
with these extreme compositions include veins and xenoliths and do not necessarily represent a significant part of
the mantle.
The EM1, EM2, and HIMU mantle reservoirs sampled
by OIB all involve crustal components that have been
introduced into the mantle (e.g. Weaver, 1991; Stracke
et al., 2005; Jackson et al., 2007). Their relatively small variation in d7Li, which contrasts with the ranges in the isotopic compositions of Sr, Nd, and Pb, is not necessarily
expected a priori and may have several explanations. (1) If
the material recycled into the mantle has d7Li values as
low as reported for some eclogites and mantle xenoliths,
the amount of Li recycled into the mantle has to be small.
In contrast, (2) if Li reintroduced into the mantle has similar d7Li values to the mantle, the isotopic composition of
Li does not provide any constraints on the amount of Li
recycled into the mantle, whereby it is irrelevant whether
the d7Li value of the reintroduced material falls in the
same range as the mantle or whether particularly low
d7Li values are balanced by material with high d7Li
values. Neither of these explanations would require a
decoupling of the Li isotopic signature from that of Sr,
Nd, and Pb, or large-scale homogenization of the Li isotopic composition in the mantle. Actually, such a homogenizationçor partial homogenization in the case of the
slightly heterogeneous OIB rocksçmay occur as late as
during melting and transport. In particular, diffusion,
which is thought to account for isotopic fractionation between wall-rock and melt (e.g. Richter et al., 2003;
Jeffcoate et al., 2007), cannot account for large-scale homogenization. At the small scale, diffusion is driven by a gradient in the chemical potential (same P and T), which
might differ for 6Li and 7Li. Diffusion minimizes the difference in chemical potential (and thus reduces the gradient) between adjacent phases (e.g. crystal and melt). This
directed material transport does not result in homogenization of the Li content or Li isotopic composition [for examples of contrasting d7Li at small scales see Brooker
et al. (2004), Vlaste¤lic et al. (2009) and Su et al. (2012)], but
eventually results in the equilibration of the chemical potential between the two phases (e.g. crystal and melt).
Once the chemical potential is the same in both phases,
continued diffusion will have the character of Brownian
motion rather than directed mass transfer and the system
will not change. In such a scenario, diffusion may explain
small-scale isotopic heterogeneities and would also imply
that such heterogeneities are long-lived [i.e. persist as long
as the system (e.g. P, T, bulk chemistry) remains unchanged] but readily adjust to changed external conditions. Diffusion, however, does not explain why pyroxenite
veins show both isotopically heavier and lighter Li isotopic
compositions than their host peridotite [see data of
Brooker et al. (2004)]. Diffusion does not result in isotopic
homogenization on the large scale (as the chemical potential gradient decreases with increasing scale) and, furthermore, operates at a rate that is considerable slower than
convective transport.
The small ranges of d7Li in different mantle reservoirs
(Fig. 5) and in carbonatites ranging in age from Archaean
to the present day overlap with the d7Li range of MORB
(Halama et al., 2008). This indicates that the Li isotopic
composition of the mantle has not changed significantly
through time (Halama et al., 2008), which in turn requires
2343
JOURNAL OF PETROLOGY
VOLUME 53
that the return of Li to the mantle is either small or isotopically not very different from the mantle. The introduction
of significant amounts of Li with an isotopic composition
that is distinctly different from the MORB source would
imply that the Li isotopic composition of the mantle becomes homogenized and that there is a temporal evolution
of d7Li in the mantle through time. Such a temporal variation in d7Li, however, is not observed (see Halama et al.,
2008). Therefore, the relation between d7Li and the isotopic composition of Sr, Nd, and Pb indicates that material
reintroduced into the mantle has on average a very similar
d7Li to that of the mantle.
CONC LUSIONS
Li contents of basalts from several OIB regions vary from
1·5 to 13·3 mg g1 and their small range in Li isotopic composition, between 2·4 and 4·8ø, largely overlaps with the
range of N-MORB. There are weak co-variations of Li
content, Li isotope composition and Sr^Nd^Pb isotope
composition of the basalts, indicating that the Li elemental
and isotopic characteristics of OIB, to some extent, reflect
mantle source heterogeneity. In comparison with the
range in Sr, Nd, and Pb isotopic composition, the variability in d7Li is small. The slightly elevated d7Li values of
HIMU-type lavas relative to average N-MORB and their
comparatively low Li contents are consistent with recycling of dehydrated oceanic crust. The enriched Li contents
of EM1 lavas along with their light Li isotope compositions
point to a sediment contribution to their mantle source
region, whereas the recycling of mantle wedge material
that has been hydrated by fluids expelled from subducting
slabs can explain the relatively high Li concentrations and
heavy d7Li values observed in EM2 basalts.
AC K N O W L E D G E M E N T S
We are very grateful to our colleagues who contributed
their technical expertise and advice to this project: C.
Schulz (Li isotopes, GFZ), U. Westernstro«er (ICP-MS,
Kiel), O. Appelt [electron microprobe analysis (EPMA),
GFZ], P. Appel (EPMA, Kiel), and G. Berger (preparation
of grain mounts for EPMA, GFZ). We thank P. Tomascak,
J. Ryan, and an anonymous reviewer for detailed comments, and D. Weis for thoughtful editorial handling.
F U N DI NG
Funding by the Deutsche Forschungsgemeinschaft through
grant GA 500/5-1 is gratefully acknowledged.
S U P P L E M E N TA RY DATA
Supplementary data for this paper are available at Journal
of Petrology online.
NUMBER 11
NOVEMBER 2012
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