ORIGINAL PAPER Deep solid-state equilibration and deep melting of plagioclase-free spinel peridotite from the slow-spreading Mid-Atlantic Ridge, ODP Leg 153 Thomas M. Will & Esther Schmädicke & Hartwig E. Frimmel Received: 10 March 2010 / Accepted: 1 September 2010 # Springer-Verlag 2010 Abstract A petrological investigation of abyssal, plagioclase-free spinel peridotite drilled during ODP cruise 153 in the North Atlantic revealed that the peridotite represent refractory, partial residual mantle material that experienced depletion of incompatible trace elements during upper mantle melting. The degree of partial melting as estimated from spinel compositions was c. 12%. Fractionated middle and heavy rare earth elements imply polybaric melting, with c. 1–4% initial melting in the garnet peridotite stability field and subsequent partial melting of ~7–10% in the spinel peridotite stability field. Geothermo- barometric investigations revealed that the solid-state equilibration of the spinel peridotite occurred at some 1,100–1,150°C and c. 20–23 kbar, corresponding to an equilibration depth of c. 70±5 km and an unusually low thermal gradient of some 11–17°C/km. A thermal re- equilibration of the peridotite occurred at ~850–1,000°C at similar depths. Naturally, the initial mantle melting in the garnet-peridotite stability field must have commenced at depths greater than 70±5 km. It is likely that the residual peridotite rose rapidly through the lithospheric cap towards the ridge axis. The exhumation of the abyssal peridotite occurred, at least in parts, via extensional detachment faulting. Given the shallow to moderate dip angles of the fault surfaces, the exhumation of the peridotite from its equilibration depth would imply an overall ridge-normal horizontal displacement of c. 50–160 km if tectonic stretching and detachment faulting were the sole exhuma- tion mechanism. Introduction Serpentinised peridotite directly exposed on the seafloor is known through dredging and drilling from several loca- tions. Examples include sections of the Mid-Atlantic Ridge (e.g. Miyashiro et al. 1969; Bonatti and Honnorez 1976; Tamura et al. 2008), the Gakkel Ridge in the Arctic Ocean (e.g. Michael et al. 2003) and the Central and Southwest Indian Ridge (e.g. Kimball et al. 1985; Dick 1989). In addition to the seafloor exposures some ophiolite com- plexes, for example, in the western Alps and Ligurian province (e.g. Lagabrielle and Cannat 1990; Rampone et al. 1995) or on the island of Zabargad in the Red Sea (e.g. Bonatti et al. 1986) provide further insights into mantle- derived peridotite that was accreted to the seafloor. Typically, these occurrences are associated with oceanic fracture zones in the close vicinity of slow-spreading ridges (with full spreading rates less than c. 5 cm/yr). Abyssal peridotite from or near slow-spreading ridge segments was recovered during several Ocean Drilling Program (ODP) expeditions (e.g. Legs 153, 209, 304/305). In this paper we focus on plagioclase-free spinel-bearing abyssal peridotite that was drilled along the Mid-Atlantic Ridge in the vicinity of the Kane Fracture Zone (‘MARK area’, 23°N; full spreading rate 2.5 cm/yr, Kong et al. 1988) during Leg 153. The primary aim of drilling Leg 153 was to characterise the variation of the mantle composition, melt migration Editorial handling: J. Raith T. M. Will (*) : H. E. Frimmel Geodynamics and Geomaterials Research Division, University of Würzburg, Am Hubland, 97074 Würzburg, Germany e-mail: [email protected]E. Schmädicke GeoZentrum Nordbayern, University of Erlangen-Nürnberg, Schlossgarten 5a, 91054 Erlangen, Germany Miner Petrol DOI 10.1007/s00710-010-0125-7
Transcript
ORIGINAL PAPER
Deep solid-state equilibration and deep meltingof plagioclase-free spinel peridotite from the slow-spreadingMid-Atlantic Ridge, ODP Leg 153
Thomas M. Will & Esther Schmädicke &
Hartwig E. Frimmel
Received: 10 March 2010 /Accepted: 1 September 2010# Springer-Verlag 2010
Abstract A petrological investigation of abyssal,plagioclase-free spinel peridotite drilled during ODP cruise153 in the North Atlantic revealed that the peridotiterepresent refractory, partial residual mantle material thatexperienced depletion of incompatible trace elementsduring upper mantle melting. The degree of partial meltingas estimated from spinel compositions was c. 12%.Fractionated middle and heavy rare earth elements implypolybaric melting, with c. 1–4% initial melting in the garnetperidotite stability field and subsequent partial melting of~7–10% in the spinel peridotite stability field. Geothermo-barometric investigations revealed that the solid-stateequilibration of the spinel peridotite occurred at some1,100–1,150°C and c. 20–23 kbar, corresponding to anequilibration depth of c. 70±5 km and an unusually lowthermal gradient of some 11–17°C/km. A thermal re-equilibration of the peridotite occurred at ~850–1,000°Cat similar depths. Naturally, the initial mantle melting in thegarnet-peridotite stability field must have commenced atdepths greater than 70±5 km. It is likely that the residualperidotite rose rapidly through the lithospheric cap towardsthe ridge axis. The exhumation of the abyssal peridotiteoccurred, at least in parts, via extensional detachment
faulting. Given the shallow to moderate dip angles of thefault surfaces, the exhumation of the peridotite from itsequilibration depth would imply an overall ridge-normalhorizontal displacement of c. 50–160 km if tectonicstretching and detachment faulting were the sole exhuma-tion mechanism.
Introduction
Serpentinised peridotite directly exposed on the seafloor isknown through dredging and drilling from several loca-tions. Examples include sections of the Mid-Atlantic Ridge(e.g. Miyashiro et al. 1969; Bonatti and Honnorez 1976;Tamura et al. 2008), the Gakkel Ridge in the Arctic Ocean(e.g. Michael et al. 2003) and the Central and SouthwestIndian Ridge (e.g. Kimball et al. 1985; Dick 1989). Inaddition to the seafloor exposures some ophiolite com-plexes, for example, in the western Alps and Ligurianprovince (e.g. Lagabrielle and Cannat 1990; Rampone et al.1995) or on the island of Zabargad in the Red Sea (e.g.Bonatti et al. 1986) provide further insights into mantle-derived peridotite that was accreted to the seafloor.Typically, these occurrences are associated with oceanicfracture zones in the close vicinity of slow-spreading ridges(with full spreading rates less than c. 5 cm/yr). Abyssalperidotite from or near slow-spreading ridge segments wasrecovered during several Ocean Drilling Program (ODP)expeditions (e.g. Legs 153, 209, 304/305). In this paper wefocus on plagioclase-free spinel-bearing abyssal peridotitethat was drilled along the Mid-Atlantic Ridge in the vicinityof the Kane Fracture Zone (‘MARK area’, 23°N; fullspreading rate 2.5 cm/yr, Kong et al. 1988) during Leg 153.
The primary aim of drilling Leg 153 was to characterisethe variation of the mantle composition, melt migration
Editorial handling: J. Raith
T. M. Will (*) :H. E. FrimmelGeodynamics and Geomaterials Research Division,University of Würzburg,Am Hubland,97074 Würzburg, Germanye-mail: [email protected]
E. SchmädickeGeoZentrum Nordbayern, University of Erlangen-Nürnberg,Schlossgarten 5a,91054 Erlangen, Germany
Miner PetrolDOI 10.1007/s00710-010-0125-7
features, hydrothermal alteration and the mode of mantledeformation. Consequently, a large amount of whole rockgeochemical (major, minor and trace elements), mineralchemical, isotope and structural data has been gathered andis summarised in Karson et al. (1997). However, to ourknowledge, rigorous geothermobarometric investigations todetermine the mantle equilibration conditions of theperidotites were not carried out. Filling this knowledgegap is the main focus of this paper. In addition, thegeochemical data presented in this paper should beconsidered as a supplement to the existing database (Karsonet al. 1997).
A schematic cross section across the MARK area(modified after Thy and Dilek 2000) is shown in Fig. 1.At site 920 of Leg 153 c. 95% variably serpentinised (~50–100%) peridotite and c. 5% gabbroic and dioritic sills,which intruded into the peridotite, were recovered from twodrill holes. Hole 920B was drilled to a depth of 126.4 mbelow sea floor (mbsf), and a depth of 200.8 mbsf wasreached in Hole 920D (Cannat et al. 1995). Core recoverywas 39.7% in Hole 920B and 47% in Hole 920D. As statedby Cannat et al. (1995) 96% of rock core recovered fromHole 920B is variably serpentinised peridotite, withporphyroclastic harzburgite forming 85% of the core. Theremaining 4% of the material recovered consists of variablyaltered lherzolite, dunite, pyroxenite, diabase, (meta)gabbroand gneissic amphibolite. Similarly, 95% of the corerecovered from Hole 920D is serpentinised peridotite.Lithologically, 83% of the peridotite is harzburgitic incomposition, the remaining rock types range from lherzoliteto dunite. The mafic rocks recovered from this hole (5%)consist of variably deformed and metamorphosedplagioclase- and pyroxene-bearing plutonic and hypabyssalrocks (Cannat et al. 1995).
As summarised by Karson et al. (1997) the recovereddrill core material has some peculiar characteristics: (1) theabundance of serpentinised mantle peridotite is very high;(2) the mantle rocks are directly overlain by a thin veneer(~20 cm) of pelagic sediments and, in some places,volcanic talus; (3) the mafic rocks (gabbro and diorite)intruded as discrete dikes and/or sills into the ultramafic
peridotite; (4) a sheeted-dike complex is missing; (5) asubmarine basaltic cover is scarce or even lacking; and(6) low-angle brittle normal faults and ductile shear zonesoccur throughout the core. Karson and Lawrence (1997)interpreted these observations in terms of an oceanic corecomplex, where the entire ‘normal’ oceanic crust, consist-ing of pillow basalt, sheeted dikes and gabbro is missing.According to these authors, highly asymmetrical, nearlyamagmatic extension along a major detachment fault led toroughly 4 km uplift of the abyssal peridotite.
Despite the serpentinisation, unaltered primary mantleminerals (olivine, orthopyroxene, clinopyroxene and spinel)are preserved in many of the peridotites. Plagioclase isnotably absent in these rocks. The presence of spinelrequires equilibration pressures of at least 8 kbar at 950°C(e.g. O’Neill 1981; Schmädicke et al. 2010). Significantlyhigher pressures in the excess of 20 kbar could be,however, a possibility. This implies an exhumation of thespinel peridotite from depths between 25 and possibly morethan 70 km, which is in strong contrast to the 4–7 km ofexhumation that have been suggested by Karson andLawrence (1997). The aim of this study is to solve thisdiscrepancy by determining the mantle equilibrium con-ditions of these rocks. This is not only important for theinference of the pressure-temperature evolution but also forthe exhumation history of the spinel peridotite.
Samples and petrography
After extensive microscopic investigation of some 150 thinsections, eight spinel peridotite samples were selected forelectron microprobe (EMP) and LA-ICP-MS analyticalwork. These samples were chosen because they containabundant fresh-looking primary minerals, summarised inTable 1.
All samples investigated show a porphyroclastic texturewith large orthopyroxene and olivine porphyroclasts. Sub-to anhedral orthopyroxene porphyroclasts are up to 12 mmin diameter and display weak deformation features such asundulose extinction, crystal bending and kinking (Fig. 2a).
Fig. 1 Schematic E-W cross-section through the MARK area(modified after Thy and Dilek2000)
Additionally, many orthopyroxene crystals contain orientedclinopyroxene lamellae (Figs. 2a, b). In places, roundedolivine grains occur at embayed orthopyroxene grainboundaries. Olivine is typically anhedral to rounded, withgrain sizes ranging from 0.1 to 4 mm, locally up to 8 mm.Some olivine crystals are recrystallised (Fig. 2c), somegrains are kinked or show undulose extinction. Porphyro-clastic anhedral clinopyroxene (cpxI) relics are up to 4 mmin diameter (Fig 2d). Locally, such porphyroclasts arereplaced by anhedral clinopyroxene II (Fig. 2e). In a fewcases, recrystallised clinopyroxene II grains (cpxII) havestraight grain boundaries and form a polygonal texture(Fig. 2f). Undulose extinction and bent clinopyroxenegrains occur locally. The 0.3–2 mm large spinel grains areyellowish to dark brown in thin section and, typically, havehighly irregular (Fig. 2 g) or vermicular, symplectite-likegrain boundaries (Fig. 2 h). Tabular shaped crystals that arelocally intergrown with clino-, orthopyroxene and/orolivine occur in places. The symplectite-like spinel grainsare locally intergrown with clinopyroxene (Fig. 2 h).Kelyphytic rims of an unidentified mineral aggregatearound spinel were observed in one sample.
Analytical methods
Mineral analyses were performed at the University ofWürzburg with a CAMECA SX 50 electron microprobe(EMP), equipped with four independent WDS spectrom-eters. Acceleration voltage was 15 kV, beam current 10 or15 nA, and counting time 20 s (30 s for Fe). Metal oxidesand synthetic mineral standards supplied by CAMECAwere used for reference and the PAP program of CAMECAfor correction procedures. The relative analytical errors areapproximately 1% for element concentrations above 2 wt.%and 5% for element abundances below <2 wt.%. Except for
spinel all iron is treated as ferrous iron. Average mineralanalyses are given in Table 2.
Trace element and rare earth element (REE) concen-trations in the minerals were determined by laser ablation(266 nm Merchantek LUV)-inductively coupled plasmamass spectrometry (Agilent 7500i) (LA-ICP-MS) at theUniversity of Erlangen. Each analysis was performed byablating spots of 50 μm in diameter at 10 Hz with anenergy density of 32–40 J/cm2 per pulse. Signal integrationtimes were 15 s for the argon background and the ablationintervals. The NIST SRM 612 glass (Pearce et al. 1997)was used for calibration. Data reduction was performedusing Si as internal standard for olivine, ortho- andclinopyroxenes and Al for spinel, based on the SiO2 andAl2O3 concentrations determined by the EMP analyses ofthe respective grains. The software program GLITTER v.3.4 (GEMOC Macquarie Research Ltd. 2000) was used fordata reduction. The obtained REE abundances are listed inTables 3 and 4. In most analyses, light REE (LREE)concentrations were below the detection limit of the LA-ICP-MS.
Results
Mineral chemistry
The compositional variations given below refer to all mineralanalyses not only to the averaged values in Table 2. The Mg-and Cr-numbers are calculated as Mg# ¼ 100 Mg=ðMgþFe2þÞ and Cr# ¼ 100 Cr=ðCr þ AlÞ, respectively.
Olivine compositions range in Mg# from 90.1 to 90.7 inall samples except for olivine grains in sample 920D 2-1,which have lower values of 88.5–89.2. The latter sample iscross-cut by a pyroxenitic vein, which might have led to theslight decrease in Mg#. Nickel concentrations in olivine
Table 1 Estimated modalproportions (vol.%) ofprimary phases (ol-olivine,opx-orthopyroxene,cpx-clinopyroxene, sp-spinel)and degree of alteration of theperidotite samples studied fromthe MARK area. mbsf—metersbelow sea floor. Interval is thedistance from the top of the coreto the core position where thesamples were taken
Deep solid-state equilibration and deep melting of plagioclase-free spinel peridotite
vary between 0.25 and 0.34 wt.% NiO in Hole 920Bsamples and from 0.27 to 0.33 wt.% in Hole 920Dperidotite. Ca concentrations vary between 266 and340 ppm. All olivine grains have very similar compositionsand are unzoned.
Orthopyroxene compositions vary little between theperidotite samples. Their Mg-numbers range from 88.9 to91.1 and the Cr-numbers vary from 10.7 to 13.5 and 11.5 to17.9 in Holes 920B and 920D, respectively. Cr2O3 rangesfrom 0.65 to 1.08 wt.% in samples from both holes, andAl2O3 varies from 3.43 to 5.45 and 2.64 to 4.49 wt.% inHoles 920B and 920D, respectively (Fig. 3a). Orthopyrox-ene grains are always more Mg-rich than coexisting olivinecrystals, indicating well-equilibrated mineral assemblages(Gurney et al. 1979). Weak intracrystalline zoning ispreserved in some large porphyroclasts, where the mineralcores are slightly higher in Cr2O3 and Al2O3 than the rims.Exsolution lamellae of clinopyroxene are present.
Clinopyroxene in all samples has high Cr contents (1.17–1.48 wt.% Cr2O3) and, as such, is classified as Cr-diopside.Na2O and TiO2 are low and generally below 0.1 wt.%. TheAl content is high (4.4–5.8 wt.% Al2O3) but fairly constantin all clinopyroxene grains (Fig. 3b). The Mg-numbers rangefrom 90.7 to 92.5 and 90.8 to 91.9 in Hole 920B and 920D,respectively (Fig. 3). The cores of clinopyroxene I haveconsistently higher Mg and lower Ca contents (19.1–20.8 wt.% MgO, 17.1–18.9 wt.% CaO) than clinopyroxene II (15.3–17.9 wt.% MgO, 20.6–23.7 wt.% CaO). The Mg-numbers ofclinopyroxene I grains are typically at the low end of Mg#range of all clinopyroxene analysed (arrows in Fig. 3b).
With one exception, the Cr# of the spinel grains is ratheruniform and varies between 26.8 and 30.6 (Fig. 4). TheMg# and the Ti content of these grains range from 64.8 to72.3 (Fig. 4a) and 0–0.1 wt.% TiO2 (Fig. 4b), respectively.One analysis in sample 920D 22-6 yielded a higher Cr# of40.6. This spinel is also characterised by the lowest Mg#determined (61.2) and the second highest Ti contentmeasured (0.09 wt.% TiO2).
Plotting the amount of the forsterite component inolivine versus the Cr# of coexisting spinel (not shown)
reveals that all samples lie within the olivine-spinel mantlearray as delineated by Arai (1994). The Cr# of coexistingspinel and clinopyroxene grains are plotted in Fig. 5.Except for the analyses from sample 920D 22-6 allperidotites are aligned along the Cr# covariation trend forspinel peridotite. This positive correlation is thought to beindicative for equilibration in the spinel peridotite field(Müntener and Manatschal 2006). The shift of the spinelCr# to the right for sample 920D 22-6 (Fig. 5) may havebeen caused by interaction with a melt phase.
Trace element compositions of clinopyroxene
The chondrite normalised trace element patterns (REEincluding Y and Sr, Zr, Ti) of all pyroxene grains analysedare similar (Figs. 6a–f) except for clinopyroxene fromsample 920D 22-6 (Fig. 6 g), which has been alreadyrecognised as being unique (Figs. 4 and 5; see above). Thelatter sample is therefore excluded from the followingdescription and will be discussed separately. The samplesare characterised by a positive slope from middle REE(MREE) to heavy REE (HREE), the chondrite-normalisedSmN range from 0.8 to 1.9, LuN from 2.8 to 8.1. In the fewsamples in which at least some LREE were above thedetection limit, the measured element concentrations indi-cate a steep positive slope from the LREE to the HREE. Allpyroxenes have very low Sr and Zr concentrations (SrN=0.02–0.07, ZrN = 0.02–0.05) and a distinct negative Tianomaly. The trace element abundances are slightly higherin the rim than in the core areas but are of similarmagnitude. The compositional variations of the clinopyr-oxene grains are rather narrow and all trace elementpatterns lie within the field typical of abyssal peridotite(Fig. 6 h). The REE pattern of clinopyroxene from sample920D 22-6 (Fig. 6 g) is almost flat and has a very gentleslope from LaN=1.3–1.4 to SmN=2.3 and LuN=3.6–8.1.Most element concentrations are slightly higher in the rimthan in the clinopyroxene core. Like the other clinopyrox-ene grains, those from sample 920D 22-6 have a strongnegative Sr (SrN=0.09–0.13) but a weak positive Zr (ZrN=2.0–2.6) anomaly. The clinopyroxene LREE concentrationsin sample 920D 22-6 are about one to two orders ofmagnitude higher than in all other clinopyroxenes, indicat-ing that the chemical composition of this sample mighthave been modified by interaction with a melt phase, whichwas already suspected on the basis of the high Cr# of spinelgrains from this sample (see above).
Pressure-temperature estimates
To our knowledge the only temperature estimates on Leg153 samples were carried out by Werner and Pilot (1997),who applied mostly fairly old calibrations of the two-
Fig. 2 Photomicrographs of spinel peridotite samples. Scale bar is0.5 mm. The holes in (d–f) are laser ablation spots. a Kinkedorthopyroxene with very narrow clinopyroxene exsolution lamellae,sample 920D 22-6, cross-polarised light. b Oriented clinopyroxenegrains in orthopyroxene host, sample 920B 12-2, cross-polarised light.c Small recrystallised olivine crystals next to olivine porphyroclastsand clinopyroxene, sample 920D 2-1, cross-polarised light. dClinopyroxene I porphyroclast core, sample 920B 13-1 (piece 1; seeTable 1), cross-polarised light. e Clinopyroxene II replacing clinopyr-oxene I, sample 920B 12-2, cross-polarised light. f Polygonal,recrystallised clinopyroxene II, sample 920B 13-1 (piece 1), plane-polarised light. g Spinel with highly irregular grain boundaries, sample920B 4-1, plane-polarised light. h Vermicular, symplectite-like spinelgrains partially intergrown with clinopyroxene in a serpentinite matrix,sample 920B 13-1 (piece 3), plane-polarised light
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Deep solid-state equilibration and deep melting of plagioclase-free spinel peridotite
Table 3 Trace elements and REE concentrations of clinopyroxene and orthopyroxene
920B4-1 cpx-c cpx-r cpx average opx-r opx-c opx-r opx-c opx average 920B12-2 cpx cpxAnalysis# 10 11 1 2 7 8 3 4Grain # A A B B C C A BType incl in opx incl in opx P P P P rx rx
cpx cpx cpx average opx-r opx-r opx average 920B13-1 cpx-r cpx-r cpx-r cpx-r cpx-c cpx cpx cpx-r6 7 8 10 10 11 12 15 16 20 21 22C C D E A B C D D E F Grx1 rx1 P P rx rx rx P P rx poly rx poly P
cpx-r cpx average opx-r opx-r opx-r opx-c opx-r opx-c opx average 920D2-1 cpx-r cpx-c cpx-r cpx-c cpx average opx-c opx-c/r opx-r opx average23 3 4 13 14 26 27 11 12 13 14 1 2 3H I I J J K K A A B B C C CP P P rx rx P P P P P P P P P
r rim analysis; c core analysis; c/r halfway between core and rim; P porphyroclast; rx recrystallised grain; rx poly recrystallised grain withpolygonal texture; bdl below detection limit.
Deep solid-state equilibration and deep melting of plagioclase-free spinel peridotite
pyroxene thermometer (Kretz 1963, 1982; Lindsley 1983;Wood and Banno 1973) and the Cr-Al in orthopyroxenethermometer of Witt-Eickschen and Seck (1991). Withoutspecifying neither the samples nor the mineral textures northe actual mineral analyses used, Werner and Pilot (1997)reported temperatures around 1,300°C, which they tenta-tively interpreted to reflect the formation of orthopyroxene-
clinopyroxene pairs. A pressure estimate was not given bythese authors, except for simply stating that the pressure didnot exceed 10 kbar.
We calculated temperatures for coexisting minerals inseveral equilibrium domains in each sample using thepressure-independent Cr-Al in orthopyroxene andorthopyroxene-olivine-spinel geothermometers of Witt-
Table 4 Trace elements and REE concentrations of clinopyroxene and orthopyroxene
920D22-2,64-68 cpx-r cpx-r cpx-r cpx average opx-r opx-c opx-r opx average 920D22-2,50-55 cpx-r cpx-c cpx-r cpx-rAnalysis# 4 5 11 8 9 10 6 7 8 9Grain # A B B C C C A A B CType rx P P P P P rx rx P P
cpx-r cpx-c cpx-c cpx average opx-c opx-r opx-r opx-c opx average 920D22-6 cpx-c cpx-r cpx average opx-r opx-c opx-r opx-c opx-r opx average10 17 18 3 4 15 16 3 4 5 6 7 8 9D E E F F G G A A B B C C Crx2 P P P P P P rx rx P P P P P
r rim analysis c core analysis; c/r halfway between core and rim; P porphyroclast; rx recrystallised grain; rx poly recrystallised grain withpolygonal texture; bdl below detection limit.
T.M. Will et al.
Eickschen and Seck (1991) and the more recent (comparedto those used by Werner and Pilot 1997) two-pyroxenethermometer calibrations of Brey and Köhler (1990) andTaylor (1998). The two-pyroxene thermometers arepressure-dependent but very robust, a 10 kbar differenceas input pressure causes only some 15–20°C difference forthe BK and c. 30°C difference for the T thermometer.For samples that have Ca contents in olivine above thedetection limit of the LA-ICP-MS, equilibrium pressureswere calculated with the Ca-in-olivine geobarometer ofKöhler and Brey (1990), using their low-temperaturecalibration. Further pressure estimations were attemptedby using the spinel Cr# in the empirical calibration ofAshchepkov et al. (2008), which is quite independent of theinput temperature (100°C difference results in up to 2 kbarpressure change). In all calculations, the results fromthe orthopyroxene-olivine-spinel thermometer of Witt-Eickschen and Seck (1991) were used as input temper-atures. The mineral assemblages used to determine theseinput temperatures involve the olivine or spinel grains
Fig. 3 Compositional variation in orthopyroxene (a) and clinopyrox-ene (b) in spinel peridotite from the MARK area. Type I clinopyrox-ene analyses are marked with arrows
0.05
0.10
0.15
0.20
0.25
0.2 0.3 0.4 0.5Cr/(Cr+Al), spinel
Cr/
(Cr+
Al),
cpx 920B4-1
920B13-1920D2-1920D22-2,50920D22-2,64920D22-6
Fig. 5 Cr# of spinel vs. Cr# of clinopyroxene in spinel peridotitefrom ODP Leg 153. Shaded field represents spinel peridotitecomposition from the Lena Through (Hellebrand and Snow 2003).All samples (except 920D 22-6) define a Cr# correlation trend; sample920D 22-6 plots slightly off-trend, possibly indicating partial (re)equilibration with plagioclase
Deep solid-state equilibration and deep melting of plagioclase-free spinel peridotite
Fig. 6 Chondrite normalised (Sun and McDonough 1989) traceelement composition of clinopyroxene in spinel peridotite. a–c Hole920B, d–g Hole 920D clinopyroxene, h summary diagram. Abyssal
peridotite field after data from Johnson et al. (1990), Johnson andDick (1992), Ross and Elthon (1997) and Tamura et al. (2008). r-rimanalysis, c-core analysis of the same grain
T.M. Will et al.
used for the Ca-in-olivine and spinel Cr# barometry,respectively.
The application of the Ca-in-olivine barometer yieldssample-averaged pressures between 19.9 and 22.9 kbar,corresponding to an equilibration depth of the spinelperidotite of some 70±5 km (Table 5). Except for sample920D 22-6, which contains spinel with the highest Cr# ofall samples (40.6), the application of the Ashchepkov et al.(2008) calibration yields identical pressures of 19–22 kbar(Table 6). These pressure estimates are consistent withphase petrological constraints, which place the garnet-spinel peridotite transition for the Cr-bearing peridotitechemical system (Cr2O3-CaO-MgO-Al2O3-SiO2) at pres-sures above 18–19 kbar (e.g. O’Neill 1981). Using thespinel Cr contents of the samples analysed, the transitionshould occur at pressures exceeding 22 kbar. Consequently,we interpret the equilibrium pressure conditions during theperidotite formation in the mantle to have been in the orderof some 20 kbar. This value is used as input for thecalculation of equilibrium temperatures with the pressure-dependent two-pyroxene geothermometers below.
The two-pyroxene thermometers yielded very consistentresults that are in good agreement with each other (Figs. 7a, c),thus pointing to equilibration between the coexisting phases.The temperature differences between the various calibrationsare generally lower than±50°C. For spinel peridotite fromHole 920B the estimated temperatures range from 900 to1,030°C (Fig. 7a). Except for some pyroxene pairs thatyielded temperatures between 1,100°C and 1,170°C, thetemperatures inferred for the remaining Hole 920D samplesare identical to the results for Hole 920B peridotites(Fig. 7c). Using the compositional data given by Niida
(1997) yields similar temperatures of 940–1,040°C (notshown) for the peridotite from Hole 920D. The pyroxenepairs that yielded high temperatures of up to 1,170°Cconsisted of clinopyroxene I with low and orthopyroxenewith high Ca contents.
Application of the Cr-Al in orthopyroxene andorthopyroxene-olivine-spinel thermometers of Witt-Eickschen and Seck (1991) gave values between 1,000°Cand 1,170°C and 1,050–1,180°C for the Hole 920B and920D samples, respectively (Figs. 7b, d). Average temper-atures obtained by these methods are around 1,100–1,150°Cand, as shown in Figs. 7b and d, correlate reasonably wellwith each other.
Discussion
The primary mantle mineral assemblage, the mineralcompositions and modal abundances in the abyssal perido-tite together with the high MgO and NiO contents in olivineare compatible with the interpretation that the rocks areresidual peridotite that experienced partial melting in theupper mantle. Even though not many LREE concentrationscould be determined, those that were measurable areconsistently very low compared to the MREE and HREE.The extreme sub-chondritic depletion of the LREE wasalready pointed out by Ross and Elthon (1997). Thus, thetrace element pattern is compatible with the refractorycharacter of the residual peridotites, which experienced adepletion of incompatible trace elements during melting inthe upper mantle. A significant refertilisation of the spinelperidotite by entrapped melts as, for example, described by
Table 5 Results from coexisting olivine-clinopyroxene barometry using the low-temperature calibration after Köhler and Brey (1990). Cacontents in olivine and clinopyroxene were determined by LA-ICP-MS measurements. The input temperatures are the average temperatures asobtained from the application of the Witt-Eickschen and Seck (1991) orthopyroxene-spinel-olivine geothermometer involving the olivine used forthe ol-cpx barometry. DCa = Caol/Cacpx. The equilibration depth is calculated using a harzburgite density of 3,300 kg/m3 (Anderson 1989)
Core Section &Piece#
Analysis# Ca,ol (ppm) Ca,cpx (ppm) lnDCa T (°C) P (kbar) P (kbar)average
Table 5 Results from coexisting olivine-clinopyroxene barometryusing the low-temperature calibration after Köhler and Brey (1990).Ca contents in olivine and clinopyroxene were determined by LA-ICP-MS measurements. The input temperatures are the averagetemperatures as obtained from the application of the Witt-Eickschen
and Seck (1991) orthopyroxene-spinel-olivine geothermometer in-volving the olivine used for the ol-cpx barometry. DCa = Caol/Cacpx.The equilibration depth is calculated using a harzburgite density of3,300 kg/m3 (Anderson 1989)
Deep solid-state equilibration and deep melting of plagioclase-free spinel peridotite
Table 6 Results from the application of the Cr-in-spinel barometer using the formulation of Ashchepkov et al. (2008). The input temperatures arethose obtained from the Witt-Eickschen and Seck (1991) orthopyroxene-spinel-olivine geothermometer involving the spinel used for theAshchepkov et al. (2008) spinel barometry. The depth estimate uses a harzburgite density of 3,300 kg/m3 (Anderson 1989). XCr, sp etc. are therespective cations in the spinel formula based on a cation sum of three for four oxygens
Core Section& Piece#
Assemblage XCr, sp XAl, sp XFe3+, sp Cr/(Cr+Al) Topx-ol-sp (°C) P (kbar) P (kbar)average
Table 6 Results from the application of the Cr-in-spinel barometerusing the formulation of Ashchepkov et al. (2008). The inputtemperatures are those obtained from the Witt-Eickschen and Seck(1991) orthopyroxene-spinel-olivine geothermometer involving the
spinel used for the Ashchepkov et al. (2008) spinel barometry. Thedepth estimate uses a harzburgite density of 3,300 kg/m3 (Anderson1989). XCr, sp etc. are the respective cations in the spinel formulabased on a cation sum of three for four oxygens
T.M. Will et al.
Tamura et al. (2008) for residual peridotite from theAtlantis Massif to the north of the MARK area, and/or a(re)-equilibration of the MARK samples in the plagioclaseperidotite stability field seems unlikely (except, possibly,for sample 920D 22-6) based on the geochemical evidenceand the thermobarometric results. This interpretation is alsoconsistent with the very low TiO2 (Fig. 4b) and Na2Oconcentrations in clinopyroxene.
As shown by Hellebrand et al. (2001) the Cr# of spineland the HREE concentrations in coexisting clinopyroxenecan be used to infer the degree of partial melting, whichwas defined as: degree of partial melting=10 ln(Cr#, sp)+24 by these authors. The correlation between these twoparameters is shown in Fig. 8. Except for sample 920D 22-6, all analyses plot on the melting trend (drawn accordingto Tamura et al. 2008) and indicate that the peridotiteexperienced c. 12% of partial melting in the upper mantle,an estimate that is consistent with earlier studies (Casey1997; Brandon et al. 2000; Alt and Shanks 2003).Hellebrand and Snow (2003) pointed out that fractionatedMREE-HREE ratios at relatively high HREE abundances inclinopyroxene in spinel peridotites from the Lena Troughcannot be explained by fractional melting in the spinelperidotite stability field only but require initial melting in
the garnet stability field. Plotting the available Leg 153 datainto the chondrite-normalised YbN in clinopyroxene vs.(Sm/Yb)N in clinopyroxene diagram of Hellebrand andSnow (2003) reveals that the residual peridotite from theMARK area can be explained by some 1–4% melting in thegarnet peridotite stability field followed by c. 7–10%melting in the spinel peridotite field (Fig. 9). Thus, initial
T(°
C)C
rAl-i
n-op
x
800
900
1000
1100
1200
800 900 1000 1100 1200
T(°C)oxp-ol-sp
(d)
920D2-1
920D22-2, 50-55
920D22-6
T(°C)2PxT
T(°
C)2
PxB
K
800 900 1000 1100 1200800
900
1000
1100
1200
(a)
920B12-2
920B4-1
920B13-1
T(°C)oxp-ol-sp
T(°
C)C
rAl-i
n-op
x
800
900
1000
1100
1200
800 900 1000 1100 1200
920B4-1
920B13-1
(b)
T(°C)2PxT
T(°
C)2
PxB
K
800 900 1000 1100 1200800
900
1000
1100
1200
(c)
920D2-1
920D22-2, 50-55
920D22-6
Fig. 7 Temperatures estimated for Hole 920B (a, b) and Hole 920D(c, d) spinel peridotite from ODP Leg 153. Two-pyroxene (2Px)thermometry was carried out using the formalisms provided by Brey
and Köhler (1990) and Taylor (1998). The Cr-Al in-opx and the oxp-ol-sp temperature estimates were determined according to thecalibrations of Witt-Eickschen and Seck (1991)
Cr/(Cr+Al), spinel
Yb
in c
px (
ppm
)
920B4-1
920B13-1
920D2-1
920D22-2,50-55
920D22-2,64-68
920D22-6
0.1
1
10
melting trend
12 % melting
8 % melting
15 % melting
0.20.10 0.3 0.4 0.5 0.6
Fig. 8 Relationship between the Cr# in spinel and the REEconcentration in clinopyroxene (expressed via the Yb content) inspinel peridotite from the MARK area. The degree of melting wasdetermined according to Hellebrand et al. (2001)
Deep solid-state equilibration and deep melting of plagioclase-free spinel peridotite
melting in the upper mantle must have started in the garnetstability field, i.e. at depths greater than the c. 75±5 km(e.g. O’Neill 1981). Moreover, continued melting in thespinel peridotite stability field must also have occurred atminimum depths of c. 70±5 km. This is the naturalconsequence of the inferred equilibration pressure of 22–23 kbar for the residual, solidified spinel peridotites.
This new finding of deep solid-state equilibration andsubsequent, presumably rapid, exhumation and accretion tothe ocean floor is supported by results recently presented byGose et al. (2009), who studied spinel peridotite from ODPLeg 153 from an entirely different angle. Using infraredspectroscopy, these authors determined the H2O contents innominally anhydrous minerals, such as orthopyroxene, andfound that the measured concentrations of 160–270 wt.ppm H2O in orthopyroxene are identical to those inoriginally deep-seated mantle xenoliths and, furthermore,are equal to, or even higher than, concentrations asdetermined by experiments at 15 kbar and 1,100°C (Rauchand Keppler 2002). Thus, the results from two conceptualvery different studies indicate much deeper spinel peridotiteequilibration conditions and exhumation from much greaterdepths than previously suggested by Karson and Lawrence(1997).
The temperatures inferred with the Witt-Eickschen andSeck (1991) thermometers and the barometric results pointto equilibration of the spinel peridotites at ~1,100–1,150°Cand 20–23 kbar, corresponding to an equilibration depth ofsome 70±5 km. A subsequent thermal equilibrationoccurred at decreasing temperatures of about 900–1,000°Cas indicated by the results of various two-pyroxenethermometers. However, there is no evidence for asubsequent low-pressure re-equilibration. This might indi-
cate that the low-temperature re-equilibration must havealso occurred at c. 70 km depth and, subsequently, that theresidual peridotites rose rapidly towards the slow-spreadingridge axis from this depth as also inferred by Gose et al.(2009) based on the H2O content in orthopyroxene. Such arelative low temperature re-equilibration together with thenear absence of ‘normal oceanic crust’ (Karson et al. 1997)seems to be consistent with exhumation of the upper mantlematerial via extensional detachment faulting in the MARKarea (e.g. Karson and Lawrence 1997). Whether theresidual peridotites were exhumed from their equilibrationdepth by pure mantle upwelling or extensional detachmentfaulting or a combination of both processes remainsunclear. Presumably, a combination of both processescaused exhumation of the residual peridotites and theiraccretion to the sub-seafloor. However, it is interesting tonote that the observed fault dips of 20–50° in the MARKarea (Karson and Lawrence 1997) imply an overallhorizontal displacement of c. 50–160 km perpendicular tothe ridge axis if the exhumation of the spinel peridotitefrom depths of some 70 km was solely accommodated bytectonic stretching and extensional detachment faulting.Remarkably, based on a seismological and seafloor mor-phology study near 13°N on the Mid-Atlantic Ridge, verysimilar magnitudes of horizontal displacement were sug-gested by Smith et al. (2006), who found evidence for thepresence of extinct core complexes at least 100 km awayfrom the active ridge axis.
Conclusions
A series of conclusions can be drawn from this study assummarised below.
Residual, plagioclase-free spinel-bearing residual mantleperidotite formed during low degrees of polybaric melting.Initial deep melting of c. 1–4% commenced in the garnet-peridotite stability field and was followed by some 7–10%melting in the spinel-peridotite stability field. Melting tookplace at depths greater than c. 70±5 km. Low degrees of‘wet’ melting (i.e. of water enriched peridotite) at mantledepths greater than 60 km was also proposed for the originof water-rich basalts from the equatorial Mid-AtlanticRidge by Ligi et al. (2005).
Solid-state equilibration of the spinel peridotite in theupper mantle occurred at ~1,100–1,150°C and approxi-mately 20–23 kbar, i.e. at c. 70±5 km depth. These deepsolid-state equilibration conditions are in agreement withrecent infrared spectroscopy study presented by Gose et al.(2009). A thermal re-equilibration at 900–1,000°C isinferred for the abyssal peridotite. The lack of re-equilibration in the plagioclase stability field probablyindicates that the spinel peridotite must have risen rapidly
1 10
1
10
0.1
YbN in cpx
(Sm
/Yb)
N in
cpx
920B 13-1
920D 22-2,50
920B 12-2
920D 22-6
920D RE 97
920B RE 97
0% g
4% g
8% g
2% sp4% sp
12% sp
8% sp
Fig. 9 Chondrite-normalised Yb vs. Sm/Yb-in-clinopyroxene dia-gram. The degree of garnet and spinel melting contours are afterHellebrand and Snow (2003). The majority of the analyses areconsistent with c. 1–4% melting in the garnet peridotite stability fieldfollowed by 7–10% melting in the spinel peridotite stability field. Thegray data points represent analyses given by Ross and Elthon (1997)
T.M. Will et al.
towards the seafloor and the ridge axis, where late-stageserpentinisation occurred. Exhumation must have beenaccompanied by cooling, which is reconciled best withextension-related detachment faulting.
Our geothermobarometric data for the Leg 153 samplesimply an unusually low geothermal gradient of 11–17°C/km (Table 6) below the Mid-Atlantic Ridge at 23 °N. Thisestimate is consistent with the very high water contents ofup to 270 ppm in orthopyroxene from residual peridotitefrom the same site (Gose et al. 2009) because, as pointedout by Ligi et al. (2005), high water and volatileconcentrations lower the mantle solidus and melting takesplace at greater depths.
Purely tectonic exhumation of the spinel-peridotitewould require a ridge-normal horizontal displacement of50 to 160 km, which is of similar magnitude as estimatesfor a Mid-Atlantic Ridge segment to the south of theMARK area (Smith et al. 2006).
Acknowledgments The research used samples provided by the OceanDrilling Program (ODP). ODP is sponsored by funding agencies of theparticipating countries under management of the Joint OceanographicInstitutions (JOI). This project was supported by a grant from theDeutsche Forschungsgemeinschaft, which is gratefully acknowledged.Peter Späthe is thanked for the superb thin section preparation, and UliSchüssler for his help with the microprobe work at the University ofWürzburg. Helene Brätz performed the LA-ICP-MS measurements at theUniversity of Erlangen. Walter Hale (ODP core repository, Bremen) isthanked not only for his assistance with sampling at the repository butalso for his great help in finding accommodation in Bremen when allhotel rooms were booked out. B. Evans and M. Okrusch are thanked fortheir reviews and J. Raith for editorial handling.
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