Inhibited eclogitization and consequences for

Gondwana Research 25 (2014) 668–684

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Inhibited eclogitization and consequences for geophysical rock properties and delamination models: Constraints from cratonic lower crustal xenoliths J. Semprich a,⁎, N.S.C. Simon b, c a b c

Physics of Geological Processes, University of Oslo, Blindern, P.O. Box 1048, 0316 Oslo, Norway Institute for Energy Technology, P.O. Box 40, 2027 Kjeller, Norway Department of Earth Sciences, University of Bergen, P.O. Box 7803, 5020 Bergen, Norway

a r t i c l e

i n f o

Article history: Received 2 April 2012 Received in revised form 6 August 2012 Accepted 28 August 2012 Available online 18 September 2012 Keywords: Cratonic xenoliths Eclogitization Metastability Rock density Dehydration melting

a b s t r a c t Studies on lower crustal and mantle xenoliths as well as geophysical data provide important information on the cratonic lithosphere. While geothermobarometric calculations of a majority of mantle xenoliths are in agreement with the typically low surface heat flow values of a craton (~40 mW/m2), P–T estimates for lower crustal xenoliths deviate significantly from the cratonic geotherms. Independent from the individual cratonic history, the temperatures are ~200–300 °C higher than what is expected at the base of the lower crust (~ 500–600 °C at ~1.3–1.6 GPa). Possible explanations may be a lack of equilibration to the cratonic geotherm or a relatively recent localized heat input. The presence of granulitic rocks under eclogite-facies conditions which are expected to prevail in the lower cratonic crust has consequences for the interpretation of geophysical rock properties. A mafic granulite which has been preserved under eclogite-facies conditions has densities and P-wave velocities similar to a felsic composition equilibrated to eclogite-facies conditions. Furthermore, phase diagrams calculated from xenolith bulk compositions demonstrate that eclogitization at relatively high temperatures as required for delamination of continental crust can only be triggered at significantly higher pressures than lithostatic at the base of the lower crust. As long as P–T conditions and the rock composition entail the assemblage to be granulitic, the addition of fluid at temperatures above 800 °C will not result in eclogitization, but rather in melt generation. This can also lead to an increase in density of up to 3%, however, this is strongly dependent on the amount of water saturation. © 2012 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction Our knowledge of the continental lower crust and the upper mantle is limited due to their inaccessibility. Even the deepest drill hole, the Kola super deep well, with a depth of ~ 12 km (e.g. Kremenetsky and Ovchinnikov, 1986; Kozlovsky, 1987; Kremenetsky et al., 1989) is far from penetrating the lower crust (located below ~ 20–25 km). To obtain information about the deep continental crust and upper mantle, researchers have to rely on geophysical methods (seismic, gravity and heat flow data) and on petrological and structural reconstruction studies of rocks interpreted to have come from the lower parts of the continental lithosphere. Studies of xenoliths and exposed terranes suggest that the lower crust mainly consists of metamorphic rocks in the granulite facies (e.g. Rudnick and Fountain, 1995) underlain by peridotitic upper mantle. Geophysical observations define the boundaries and physical properties of both the lower crust and the lithospheric upper mantle. Furthermore the seismic velocity structure in different tectonic settings is reasonably well known (e.g. Holbrook et al., 1992; Christensen and Mooney, 1995; Rudnick ⁎ Corresponding author. Tel.: +47 22 85 69 21; fax: +47 22 85 51 01. E-mail address: [email protected] (J. Semprich).

and Fountain, 1995; Sapin and Hirn, 1997). However, the lateral resolution of wide-angle soundings is relatively poor resulting in averaged velocity data over large areas. The velocity layers therefore may be composed of combinations of rock types (e.g. Kuusisto et al., 2006). Furthermore, the scale and spatial distribution of heterogeneities can result in an observed seismic velocity that differs from the true average velocity of the layer (e.g. Brittan and Warner, 1996, 1997). Since a variety of rock type mixtures can represent an observed seismic velocity, the petrological nature of the lower crust and the extent of equilibration to lower crustal conditions cannot be determined unambiguously on the basis of seismic velocity data alone (e.g. Downes, 1993; Rudnick and Fountain, 1995; Kuusisto et al., 2006). Information about the mineral assemblage, chemical composition and geochronology, however, can only be obtained by direct sampling of lower crustal rocks. In contrast to granulite terranes, these crustal and mantle xenoliths are regarded as far less retrogressed and altered since they are brought to the surface by rapid magmatic eruptions. Therefore, they are often interpreted to record the mineral chemistry and equilibria of the lower crust and upper mantle (e.g. Griffin et al., 1979; O'Reilly and Griffin, 1985; Lu et al., 2012; Xu et al., 2012). Since lower crustal xenoliths can be difficult to recognize, three criteria have been suggested: 1) the samples should differ

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J. Semprich, N.S.C. Simon / Gondwana Research 25 (2014) 668–684

from nearby outcropping rocks; 2) P–T estimates should coincide with inferred lower crustal conditions; 3) internal mineral isochrons should give eruption ages (Griffin and O'Reilly, 1987). However, in tectonically stable regions with thickened crust and low steady state surface heat flow (e.g. the Fennoscandian Shield), modeled geotherms which are in agreement with a majority of garnet-bearing mantle peridotites (e.g. Kukkonen and Jõeleht, 1996; Kukkonen and Peltonen, 1999) predict eclogite-facies conditions for the lower crust (e.g. Peltonen et al., 2006; Kukkonen et al., 2008). In contrast, the overwhelming majority of lower crustal xenoliths show granulite-facies assemblages and yield P–T estimates compatible with high-pressure granulite-facies conditions (Kempton et al., 1995, 2001; Hölttä et al., 2000; Peltonen et al., 2006). Thus, there is a significant discrepancy between the geophysical derived lower crustal P–T conditions and the calculated estimates based on mineral chemistry (ΔT of ~ 200– 300 °C at P of ~ 1.3–1.6 GPa). This discrepancy is commonly explained by the preservation of lower crustal rocks in a metastable condition (e.g. Peltonen et al., 2006; Kukkonen et al., 2008). The P–T estimates calculated from granulite xenoliths are therefore not interpreted as in-situ conditions in the crust but rather as preserved conditions of earlier metamorphic or magmatic events (Downes, 1993; Kempton et al., 1995, 2001; Pearson et al., 1995; Hölttä et al., 2000; Schmitz and Bowring, 2000; Peltonen et al., 2006; Kukkonen et al., 2008). Furthermore, most xenoliths are older than their host diatremes (e.g. Rudnick and Fountain, 1995), with most xenolith isochron ages reflecting the preceding regional metamorphic and magmatic history of the area (Kempton et al., 1995, 2001; Hölttä et al., 2000; Schmitz and Bowring, 2000; Peltonen et al., 2006). In the first part of this paper we will compare previously published P–T estimates of lower crustal and mantle xenoliths from selected cratonic regions (Fennoscandian Shield and Kaapvaal craton) with surface heat flow geotherms and evaluate if and which portions of the lower crust and upper mantle could be in a metastable condition. In the second part, we will quantify the effects of compositional variations and metastability on physical rock properties such as densities and seismic velocities using the measured rock compositions of lower crustal xenoliths as input. Furthermore, we will discuss the potential for eclogitization and subsequent delamination of the lower crust which has been proposed for example for the Fennoscandian Shield (Kukkonen et al., 2008). A discussion of the effects of fluids and fluid-induced melting will conclude the paper. 2. Cratonic lower crustal and mantle xenolith data compared to thermal models 2.1. Characteristic cratonic features The transformation of early continental lithosphere into cratons requires the development of tectonically stable regions with thickened crust and low steady state surface heat flow with a conductive geotherm of ~ 40 mW/m 2 (e.g. Pollack and Chapman, 1977). This apparent stability over geological timescales as well as a significant thickness of the crust (~40–60 km) characterizes a craton. In most Archean cratons, the lithosphere underwent a few cycles of extension and rifting while amalgamating smaller terranes to the final craton. At the end of the Archean, crustal heat production was approximately double the present value (Michaut et al., 2009). Despite the elevated temperatures in some cratons during the Archean, the root was sufficiently cold to remain stable and to allow the preservation of rocks of Archean age (e.g. Mareschal and Jaupart, 2006; Michaut et al., 2009). In addition, Archean cratons are characteristically underlain by thick, high-velocity roots extending depths of ~ 225 km (e.g. Sleep, 2003 and references therein), whereas the adjacent mobile belts lack evidence for similar structures (e.g. James et al., 2001). For this study, we have reviewed published data for selected suites of xenoliths from the cratonic lower crust and upper mantle. In most

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cases the latest metamorphic or magmatic event prior to the magmatic activity which was responsible for xenolith entrainment dates back to approximately a billion years. Therefore, we can assume that the xenoliths have been located in the lower crust for a significant amount of time. This allows us to draw conclusions about the time scale at which thermal and/or chemical equilibration does or does not happen. The two representative cratonic areas we have chosen for this paper not only show similarities but also differences. The crust beneath the Fennoscandian Shield is unusually thick (up to 60–63 km at the Archean-Proterozoic boundary zone in eastern Finland) but thins out towards the core of the craton to 40–45 km (Peltonen et al., 2006; Kukkonen et al., 2008). In contrast, the lower crust within the Kaapvaal craton is represented by only ~ 15 km of intermediate-velocity rocks and the depth to the Moho is ~ 37 km (e.g. James et al., 2003). The Fennoscandian Shield is characterized by a low surface heat flow value of 36 mW/m2 (Kukkonen and Peltonen, 1999). The surface heat flow of the Kaapvaal craton is with ~ 45 mW/m 2 somewhat higher but still low compared to the adjacent mobile belts with values as high as 80 mW/m 2 and an estimated lithospheric thickness of only ~ 120 km in contrast to more than 200 km beneath the craton. Fig. 1. shows two reference geotherms based on the surface heat flow, the distribution of heat production sources and the variation of thermal conductivity with temperature (Pollack and Chapman, 1977). The first represents a 40 mW/m2 standard cratonic heat flow geotherm while the second is typical for an average rift geotherm of 90 mW/m 2 (Pollack and Chapman, 1977). Both are completely independent from P–T data obtained from mantle xenoliths. In addition, a mantle xenolith calibrated geotherm derived from a lithospheric heat transfer model for the Fennoscandian Shield (Kukkonen and Jõeleht, 1996; Kukkonen and Peltonen, 1999) and a xenolith-derived geotherm for the Kalahari (Kaapvaal and Zimbabwe) craton (Rudnick and Nyblade, 1999) are presented in Fig. 1. 2.2. Mantle and lower crustal xenolith data A small number of peridotite xenoliths have been reported from the Kola Peninsula and have been classified as spinel lherzolites and wehrlites (Beard et al., 2007; mineral assemblages listed in Table 1). Temperature estimates at an assigned pressure of 1.0 to 1.5 GPa as calculated by Beard et al. (2007) are given in Table 1 together with the applied thermometers. The lower crustal xenoliths from the Kola Peninsula (Belomorian Province) vary from feldspar-poor to feldspar-rich mafic garnet granulites (Kempton et al., 1995, 2001; Table 1). The feldspar-poor granulites have been described as “eclogitic” but, by definition, eclogites do not contain any plagioclase which is why we refer to these rocks as “Pl-poor granulites”. In addition, quartz-rich, felsic (migmatitic) granulites as well as pyroxenites and composite xenoliths have been categorized (Kempton et al., 1995). These xenoliths yield P–T estimates of 750–930 °C and 1.2–1.5 GPa (Kempton et al., 1995; Table 1; Fig. 1). The mantle beneath the Karelian craton is represented by garnet– spinel peridotites, garnet lherzolites, harzburgites, wehrlites and some websterites (Peltonen et al., 1999; Lehtonen et al., 2004; Peltonen and Brügmann, 2006; Table 1). Eclogites (some diamondiferous) constitute a small portion of the xenoliths and are interpreted to represent mantle-derived melts or cumulates rather than subducted oceanic lithosphere (Peltonen et al., 1999, 2002). Fig. 1 shows P–T estimates of garnet and garnet–spinel mantle xenoliths calculated by Kukkonen and Peltonen (1999). The Karelian lower crustal xenoliths are predominantly mafic granulites with high abundances of amphibole and biotite (Hölttä et al., 2000; Peltonen et al., 2006; Table 1). Minor amounts of orthopyroxene-bearing gabbros and felsic granulites are also present in the xenolith suite. Thermobarometry yields T of 750–925 °C at 0.75–1.25 GPa (Hölttä et al., 2000; Table 1; Fig. 1).

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Fig. 1. Plot of P–T estimates of lower crustal and mantle xenoliths from Kola, Karelia and Kaapvaal. Sources for each data set are listed in Table 1. The data labeled as “Kaapvaal mantle S” are results from Simon et al. (2003, 2007). “Kaapvaal mantle J” are results given in James et al. (2004). A standard cratonic heat flow geotherm is given as 40 mW/m2 while an average rift geotherm is represented by 90 mW/m2 (Pollack and Chapman, 1977). “FS” represents the xenolith-controlled heat-flow geotherm for the central Fennoscandian shield (Kukkonen and Peltonen, 1999). The xenolith based geotherm for the Kaapvaal craton constructed by Rudnick and Nyblade (1999) is given as “K”.

Mantle xenoliths were entrained in a number of kimberlite pipes erupting through the Kaapvaal craton and its margins and can be divided into spinel lherzolites and harzburgite, low-T garnet lherzolites and harzburgites and high-T and sheared garnet lherzolites (James et al., 2004). A majority of the low-T garnet peridotites from Kimberley and Northern Lesotho yield T between 900 and 1100 °C at 3.0– 4.5 GPa (Simon et al., 2003, 2007; Table 1; Fig. 1). James et al. (2004) reported T between 800 and 1200 °C at 3.1–5.8 GPa (Table 1; Fig. 1). Graphite- and diamond-bearing eclogite xenoliths of mantle origin have also been reported (e.g. Roberts Victor and Bellsbank, where they dominate the xenolith suite). Equilibration temperatures for most of the mantle eclogites are ~ 1150 °C at an assigned P of 5.0 GPa (Viljoen, 1995; Table 1). The lower crustal mafic xenolith suite from the southwestern margin of the Kaapvaal craton consists of garnet pyroxenites, eclogites and mafic granulites (Pearson et al., 1995; Table 1). A comparison of lower crustal xenoliths from the southwestern margin of the Kaapvaal craton entrained in “on-craton” and “off-craton” kimberlites yields two distinct P–T arrays (Pearson et al., 1995). Granulite xenoliths found in kimberlites away from the craton margin (“off-craton”) define a trend from 675 °C, 0.7 GPa to 850 °C, 1.5 GPa while samples from kimberlites near the margin and on the craton produce a similar trend but displaced towards lower temperatures by approximately 100 °C (Pearson et al., 1995; Table 1; Fig. 1). There seems to be a paucity of granulite xenoliths from the central part of the craton except for ultra-high temperature (UHT) granulites from the Free State kimberlites comprising varying proportions of quartz, garnet, sapphirine, sillimanite, plagioclase and accessories (Dawson and Smith, 1987; Dawson et al., 1997). The granulites record P–T conditions of 900–1000 °C and 0.9–1.1 GPa (Dawson and Smith, 1987; Dawson et al., 1997; Schmitz and Bowring, 2003b). 2.3. Mantle xenolith data compared to cratonic geotherms and assessment of equilibration In this section, we review P–T estimates from geothermobarometry of spinel and garnet peridotite xenoliths and compare these values to

geotherms derived from geophysical data and geodynamic considerations. A major drawback of spinel peridotites is the lack of a reliable barometer making the calculation of temperatures only possible at assigned pressure estimates. Pressures are usually chosen by projection onto the cratonic geotherm and restricted by the spinel stability field. Therefore deviations of spinel peridotites from cratonic geotherms have been reported repeatedly (Boyd et al., 1999; Simon et al., 2003, 2007; Beard et al., 2007; Table 1). This discrepancy may be due to an underestimation of the pressure in these samples since the spinel stability field also depends on the mantle composition shifting the spinel–garnet phase transition in refractory peridotite to higher pressure (e.g. Basu and MacGregor, 1975; O'Neill, 1981). The possibility of a metastable overlap of the spinel and garnet stability fields has also been proposed (e.g. Boyd et al., 1999). Furthermore some samples may represent highly depleted and therefore Al-poor harzburgites equilibrated in the garnet stability field without detectable amounts of either garnet or spinel (e.g. Simon et al., 2003). Alternatively spinel may have formed by garnet break down at a very late stage in the xenolith history as indicated by spinel–pyroxene symplectites interpreted as pseudomorphs after garnet (e.g. Simon et al., 2003). If all of the above can be excluded, there is a possibility that the temperature estimates deviating from the geotherm may reflect incomplete major element (especially Al and Cr) re-equilibration during cooling and thus do not reflect the actual temperature at the time of entrainment since the spinel peridotite xenoliths originate from shallow levels (~60 km) where the ambient temperatures (~600–700 °C) are too low for complete chemical equilibration (Boyd et al., 1999). Alternatively, the unusually high temperature estimates may be real and reflect relatively recent additional heat input at the crust–mantle boundary (e.g. magmatism, regional metamorphism, shear heating) and the lack of thermal re-equilibration to lower ambient conditions due to insufficient time. For example, the observed high temperatures in spinel-bearing xenoliths from the Kaapvaal craton, (Simon et al., 2007) are consistent with a heating event at ~1.0 Ga recorded in lower crustal xenoliths from northern Lesotho (Schmitz and Bowring, 2003a). This implies that the lower crust and uppermost part of the mantle have been preserved in a metastable condition either due to lack of time or as a result of insufficiently

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Table 1 Geochronology of diatremes and xenoliths, lithologies, geothermobarometry and data sources for the Fennoscandian Shield and the Kaapvaal craton. Type and age of host rock and dating method Fennoscandian Shield Monticellite kimberlite (Kandalaksha): 382 ± 14 Ma 365 ± 16 Ma K-Ar (Beard et al., 1998)

Lamprophyres (Elovy Island): 368 ± 15 Ma 360 ± 16 Ma K–Ar (Beard et al., 1996)

Rock type mineralogy

Age of xenoliths and dating method

P–T estimates and applied thermo-/barometer

Spinel lherzolites: Ol + Opx + Cpx + Spl ± (Ti-rich and Ti-poor) Phl ± (Ti-rich and Ti-poor) Amp ± Ap ± Ilma (Beard et al., 2007)

Not available

Wehrlites: Ol + Cpx + Spl ± (Ti-rich and Ti-poor) Phl ± (Ti-rich and Ti-poor) Amp ± Ap ± Ilm ± (relict) Opx (Beard et al., 2007) Mafic granulites: Cpx + Grt + Pl (5–85 vol.%) + Rt ± Qtz ± Opx ± Phl (≤45 vol.%) ± Amp ± Scp± Ttn ± Zrn (K., 1995b; K., 2001c)

Not available

T: 775–969 °C two-pyroxene (Brey and Köhler, 1990) two-pyroxene (Wells, 1977) Ca/Al-in-Opx (Witt-Eickschen and Seck, 1991) P: assigned to 1.0–1.5 GPa T: 817–904 °C Ca-in-Opx (Brey and Köhler, 1990) Ca/Al-in-Opx (Witt-Eickschen and Seck, 1991) based on rare Opx only P: assigned to 1.0–1.5 GPa T: 750 ± 25 °C Grt–Cpx (Powell, 1985) P: 1.2–1.5 GPa Grt + Cpx + Pl+ Qtz (Newton and Perkins, 1982)

Migmatitic granulites: Grt + Cpx ± Pl ± Qtz ± Zrn (mesosome) Pl + Kfs + Qtz (leucosome) (K., 1995; K., 2001)

Kimberlites (Lahtojoki): 560–430 Ma K–Ar (Tyni, 1997)

Kaapvaal Craton Kimberlites (Kimberley): ~200–110 Ma

Pyroxenites: Cpx ± Phl ± Amp (Kempton et al., 1995; Kempton et al., 2001) Mantle, upper layer (~60–~110 km): Garnet–spinel harzburgites Ol + Opx + Grt + Spl ± Cpx Middle layer (110–180 km): garnet harzburgites, lherzolites, wehrlites and websterites Lowest layer (180–250 km): Garnet lherzolites (Peltonen et al., 1999; Lehtonen et al., 2004; Peltonen and Brügmann, 2006) Mantle eclogites: Grt + Cpx ± Dia ± Gr (Peltonen et al., 2002)

2.84 Ga, U–Pb, Zrn (D., 2002d) 2.8–2.0 Ga, Nd model ages (K., 2001) 2.59 Ga, U–Pb, Grt (K., 2001) 2.56 ± 0.41 Ga, Pb–Pb, sample (K., 2001) 2.47 Ga, U–Pb, Zrn (D., 2002) 2.23 ± 0.47 Ga, Pb–Pb, WR (K., 2001) 2.12 ± 0.11 Ga, Pb–Pb, sample (K., 2001) 2.10 ± 0.10 Ga, Ar–Ar, Phl core (K., 2001) ~1.9 Ga, Rb–Sr, WR (K., 2001) 1.86 ± 0.51 Ga, Pb–Pb, sample (K., 2001) 1.83 Ga, U–Pb, Zrn (K., 2001) 1.77–1.61 Ga, U–Pb, Zrn (D., 2002) ~1.5 Ga, Sm–Nd isochrons (K., 2001) 1.47 and 1.45 Ga, U–Pb, Zrn (D., 2002) 1.97 ± 0.46 Ga, Pb–Pb, migmatitic granulite (K., 2001)

394 ± 2 Ma, Ar–Ar, Hbl, (K., 2001)

T: 930 °C Grt–Cpx (Powell, 1985) P: 1.5 GPa Grt + Cpx + Pl+ Qtz (Newton and Perkins, 1982) Not available

~3.3 Ga, Re–Os model age of middle layer (Peltonen and Brügmann, 2006)

T: 1100–1400 °C Two-pyroxene (Brey and Köhler, 1990) P: 3.3–7.2 GPa Al-in-Opx (Brey and Köhler, 1990)

Not available

T: 1117–1500 °C Grt–Cpx (Ellis and Green, 1979) P: 4.9–6.1 GPa Ca-Tschermaks in Cpx (Simakov and Taylor, 2000) T: 750–925 °C TWEEQU (Berman, 1991) P: 0.75–1.25 GPa TWEEQU (Berman, 1991) (corrected to) 0.9-1.4 GPa (Peltonen et al., 2006)

Mafic granulites: Cpx + Amp (≤60%) + Pl ± Grt ± Opx ± Bt ± Qtz (Hölttä et al., 2000; Peltonen et al., 2006)

3.53 ± 0.10 Ga, U–Pb, Zrn (Peltonen et al., 2006) ~3.7 Ga, Nd model age (Peltonen et al., 2006) 3.47 ± 0.10 Ga; 3.48 ± 0.10 Ga, U–Pb, Zrn (Peltonen et al., 2006) ~3.37 Ga; 3.2–2.9 Ga, Nd model age (Peltonen et al., 2006) 2.3–2.1 Ga, Nd model ages (Peltonen et al., 2006) 2.0–1.7 Ga,U–Pb, Zrn (Hölttä et al., 2000; Peltonen et al., 2006) ~1.9–1.6 Ga, Pb–Pb, Grt–Pl (Peltonen et al., 2006) 1.72; 1.69; 1.6 Ga, Sm–Nd, Grt–Cpx (Hölttä et al., 2000; Peltonen et al., 2006) 498 ± 7; 459 ± 19 Ma, Rb–Sr, Cpx–Pl (Peltonen et al., 2006)

Mantle xenoliths: Spinel lherzolites and harzburgites (b100 km) low-T garnet lherzolites and harzburgites

3.2 Ga, Re–Os (Simon et al., 2007)

Spinel peridotites: T: 945–1100 °C Ol–Spl (Ballhaus et al., 1991) (continued on next page)

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Table 1 (continued) Type and age of host rock and dating method Northern Lesotho: 95 Ma–87 Ma

Kimberlites (e.g. Bellsbank): ~118 Ma (Viljoen, 1995 and references therein) Kimberlites (Namaqua-Natal): 143–74 Ma Northern Lesotho: 95 Ma–87 Ma (Schmitz and Bowring, 2003a and references therein)

Kimberlites (Free State): 132–88 Ma (Schmitz and Bowring, 2003a and references therein) a b c d e f

Rock type mineralogy

Age of xenoliths and dating method

(100–180 km); high-T garnet lherzolites (>175 km) (Simon et al., 2003; James et al., 2004; Simon et al., 2007)

P–T estimates and applied thermo-/barometer P: assigned to 2.0 GPa Low-T Grt peridotites : T: 900–1100 °C (S., 2007 e) T: 804–1207 °C (J., 2004 f) Grt–Ol (O'Neill and Wood, 1979) P: 3.0–4.5 GPa (S., 2007) Al-in-Opx (Brey and Köhler, 1990) P: 3.1–5.8 GPa (J., 2004) Al-in-Opx (MacGregor, 1974) High-T Grt peridotites (J., 2004): T: 1122–1570 °C Grt–Ol (O'Neill and Wood, 1979) P: 5.1–7.6 GPa Al-in Opx (MacGregor, 1974) T: ~1150 °C Grt–Cpx (Ellis and Green, 1979) P: assigned to 5.0

Mantle eclogites: Grt + Cpx ± Dia ± Gr

2.9 Ga Re–Os isochron, (Richardson et al., 2001)

Garnet pyroxenites: Grt + Cpx + Opx ± Spl (Pearson et al., 1995) Garnet granulites: Cpx + Grt + Pl + Qtz ± Amp Two pyroxene garnet granulites: Cpx + Opx + Grt + Pl ± Amp Kyanite granulites: Cpx + Grt + Ky + Pl + Qtz ± Opx (Pearson et al., 1995)

Not available

T: 702–896 °C Grt–Cpx (Ellis and Green, 1979)

3.6–1.3 Ga; Nd model ages; Markt (Huang et al., 1995) 2.0–1.0 Ga; Nd model ages; Lesotho (Huang et al., 1995) 2.93 ± 0.07 and 2.91 ± 0.08 Ga; Sm–Nd; Grt–Cpx–Pl–WR (Huang et al., 1995) 1.49 ± 0.36 Ga; Pb–Pb, WR (Huang et al., 1995) 1.24 ± 0.17 Ga and 1.18 ± 0.20 Ga, Sm–Nd, WR (Huang et al., 1995) 1.1–1.0 Ga, U–Pb, Zrn and Mnz (Schmitz and Bowring, 2000; 2003) 851 ± 56 Ma; 604 ± 36 Ma and 595 ± 10 Ma; Sm–Nd; Grt–Cpx–Pl–WR (Huang et al., 1995) 2.63 Ga, U–Pb, Zrn (Schmitz and Bowring, 2000)

T: 675–850 °C (“off-craton”) T: 608–825 °C (“on-craton”) Grt-Cpx (Ellis and Green, 1979) P: 0.7–1.5 GPa (“off-craton”) P: 0.6–1.7 (“on-craton”) Grt + Cpx + Pl+ Qtz (Newton and Perkins, 1982)

Eclogites: Grt + Cpx + Qtz ± Ky ± Rt (Pearson et al., 1995) UHT granulite xenoliths: Grt + Qtz + Srp + Sil ± Opx ± Ky ± Crn ± Kfs (Dawson and Smith, 1987, Dawson et al., 1997)

2.72–2.71 Ga, U–Pb, Zrn and Mnz (Schmitz and Bowring, 2003a, 2003b)

T: 760–1050 °C Grt–Cpx (Ellis and Green, 1979) P: assigned to 1.5 GPa T:900–1000 °C P: 0.9–1.1 GPa (petrogenetic grid constraints; Dawson et al., 1997)

All mineral abbreviations after Whitney and Evans (2010). Kempton et al. (1995). Kempton et al. (2001). Downes et al. (2002). Simon et al. (2007). James et al. (2004).

high T for chemical re-equilibration. Since a time span of ~0.9 Ga between this last regional heating event (~1.0 Ga) and the kimberlite eruption (~80–100 Ma) should be sufficient for a large-scale thermal re-equilibration and major-element equilibration on the grain- to thin section scale at elevated T exceeding 800 °C (e.g. element diffusion within several ten thousand years; Griffin et al., 1999) inhibited equilibration due to low ambient T (b700 °C) is more likely in this case (unless there has been a younger event which could not be detected by geochronology). In addition, disequilibrium textures and metasomatic alteration have been described in many spinel peridotite xenoliths (Boyd et al., 1999; Beard et al., 2007). Certain metasomatic features have been attributed to the interaction of the kimberlitic magma with the xenoliths during or shortly before the eruption such as the crystallization of secondary clinopyroxene (Simon et al., 2003, 2007; Beard et al., 2007) while others seem to be older than the magmatism such as the Ti-rich and Ti-poor minerals found in the xenoliths from Kola (Beard et al., 2007). However, it is not clear to which extent this metasomatism may have influenced the major element chemistry of the primary minerals and therefore the temperature estimates of these xenoliths.

In contrast, the majority of low-T garnet-bearing peridotites from Karelia and Kaapvaal reflect cold mantle conditions and plot on P–T trends consistent with the low surface heat flow geotherm of 40 mW/m 2 which is also reasonably close to the xenolith-based geotherms for both cratons (Fig. 1). Two arrays for the mantle xenoliths of the Kaapvaal craton are shown in Fig. 1 because a different combination of thermobarometers has been chosen as best fit by the respective authors (James et al., 2004; Simon et al., 2007; Table 1). Despite evidence for metasomatism of at least parts of the upper mantle in both cratons and for kimberlite modification of the peridotites (Simon et al., 2003, 2007; Peltonen and Brügmann, 2006 and references therein), the mineral major element data appear to record coherent P–T information. This suggests that there was enough time at sufficient temperatures ensuring major element equilibration between phases before xenolith entrainment and a relatively rapid major element diffusion after kimberlite modification on which the thermometers are based (Simon et al., 2003, 2007). These findings can still be verified when a numerical approach to paleogeotherm fitting and a larger database including texturally unequilibrated samples are used (Mather et al., 2011).

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Fig. 2. A comparison of geochronology data from Kola, Karelia and Kaapvaal with regional geological events showing the wide scatter of ages throughout the Archean and Proterozoic. Data sources are given in Table 1. Abbreviations: Zrn = zircon; Grt = garnet; Phl = phlogopite; Hbl = hornblende; Cpx = clinopyroxene; Pl = plagioclase; s = individual sample; WR = whole rock. Interpretation on the right side of each column according to Kempton et al. (1995, 2001) for Kola, Hölttä et al. (2000) and Peltonen et al. (2006) for Karelia and Schmitz and Bowring (2000, 2003a, 2003b) for Kaapvaal.

There is also evidence for significant metasomatism in the high-T garnet peridotites and it has been suggested that they may have preserved chemical disequilibrium in certain components making them unsuitable for thermobarometry (James et al., 2004). Alternatively, the high temperatures may reflect a recent thermal and chemical perturbation close to the base of the mantle lithosphere caused by the transition from a conductive to convective geotherm. Smith and Boyd (1987) have associated the high-T metasomatism (1200–1400 °C) with the infiltration of mafic melts resulting in the formation of significant amounts of garnet and clinopyroxene shortly before xenolith entrainment (see also Griffin et al., 2009 and references therein). Despite the complications discussed in this chapter, we conclude that the majority of cratonic mantle xenolith P–T data correlate well with a cold conductive geotherm as demonstrated in Fig. 1. Deviations from the low surface heat flow geotherm have not only been recorded in the uppermost part of the mantle where the lack of equilibration may be the result of insufficient temperatures, chemical disequilibrium or metasomatism but also in regions with very high temperatures where insufficient time for the re-equilibration to ambient P–T conditions could be a major factor. 2.4. Lower crustal xenolith data compared to cratonic geotherms and assessment of equilibration Geothermobarometry of lower crustal granulites from the cratonic areas of Kola, Karelia and Kaapvaal yields much higher T

estimates (Fig. 1) than what is expected for old and thick Archean crust and is incompatible with the 40 mW/m 2 surface heat flow geotherm. Instead, the P–T arrays correspond well to an average rift geotherm of 90 mW/m 2 (Pollack and Chapman, 1977; Fig. 1). The T estimates mainly range between 700 and 800 °C and are therefore ~ 200–300 °C higher than the assumed values at the base of the lower cratonic crust (~ 500–600 °C). This feature is independent of the regional history of the craton, the relative location of the diatremes in which the xenoliths are found (center or margin), the applied thermometers and barometers and the rock type (mafic or more felsic). Therefore, the P–T data derived from granulite xenoliths are mostly interpreted by most authors as reflecting conditions of metamorphism attained during the most recent orogeny or heating by the latest magmatic event before the eruption of the diatreme, rather than representing actual conditions in the lower crust at the time of xenolith entrainment (Pearson et al., 1995; Hölttä et al., 2000; Schmitz and Bowring, 2000; Kempton et al., 2001; Peltonen et al., 2006). This is also supported by isotopic ages which coincide with the older regional orogenic and magmatic events such as the Svecofennian orogeny of the Fennoscandian Shield (Kempton et al., 1995, 2001; Hölttä et al., 2000; Peltonen et al., 2006; Fig. 2), the Namaqua orogeny at the southwestern margin of the Kaapvaal craton (Schmitz and Bowring, 2000; Fig. 2), and the Ventersdorp flood basalt which may have influenced the UHT granulites from the Free State kimberlites in the center of the Kaapvaal craton (Schmitz and Bowring, 2003b; Fig. 2). Since a time span of ≥ 1.0 Ga appears to be

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sufficient for a major element equilibration on the grain scale several workers (e.g. Hölttä et al., 2000; Peltonen et al., 2006; Kukkonen et al., 2008) argued that the high grade rocks are preserved due to lack of re-equilibration to lower P–T conditions because element diffusion velocities in minerals decrease with decreasing temperature (D= D0e -E/RT for thermally activated diffusion, e.g. Dodson, 1973). Many xenoliths show evidence for one or more episodes of metasomatic alteration either in form of biotite or amphibole (Kempton et al., 1995, 2001; Hölttä et al., 2000; Peltonen et al., 2006). While the phlogopite cores from the Kola Peninsula were dated and may be closely associated with regional metamorphic and magmatic events (Kelley and Wartho, 2000; Kempton et al., 2001; Fig. 2), the formation of amphiboles in the Karelian samples lack time constraints. However, some textural features such as the replacement of garnet by amphibole (Hölttä et al., 2000) suggest that at least some amphiboles are secondary and have formed at a late stage in the metamorphic history of the samples. Nonetheless, the coexistence of completely dry granulites and partly metasomatized samples with relict granulitic assemblages in the same xenolith suite indicates that the samples either originally contained various amounts of hydrous phases (i.e. crystallization of some samples from volatile-enriched residues; Hölttä et al., 2000), or have been hydrated to varying degrees during metamorphic processes (Hölttä et al., 2000; Kempton et al., 2001). The first alternative should have resulted in textural equilibration of the hydrous phases with the other minerals as well as similar ages. As this has not been reported in the reviewed literature (Kempton et al., 1995, 2001; Hölttä et al., 2000), we propose that the majority of fluids have infiltrated the samples at a later stage. This suggests that retrogression is only possible with the addition of fluids while dry high-grade rocks will not re-equilibrate to changing (lower) T conditions in a stable tectonic environment (e.g. Austrheim, 1987; Glodny et al., 2003, 2008). In summary, there is ample evidence for fluid infiltration and alteration of lower crustal rocks but no xenolith with reported eclogite facies mineral assemblage or signs of partial eclogitization, despite the cratonic geotherm predicting eclogite-facies conditions in the lower crust (Kukkonen and Peltonen, 1999). We cannot exclude the possibility of selective sampling by the kimberlite or the delamination of all eclogitized lower crust but it seems unlikely that a heterogeneous layer with highly variable amounts of water-bearing phases (as indicated by the xenolith evidence) should be eclogitized and delaminated uniformly. Hence, we propose that the fluid infiltrated the rocks shortly after the magmatism responsible for the formation of the Proterozoic granulites when the temperatures in the lower crust were still too high for eclogite-facies conditions. In contrast, the xenolith suite from the margin of the Kaapvaal craton contains a significant amount of eclogitic samples which have been interpreted as the result of a continuous transition from granulite to eclogite-facies conditions and related equilibration of the mineral assemblage (Pearson et al., 1995). However, zircons in some of the eclogite xenoliths are significantly older than the granulite-facies metamorphism (Schmitz and Bowring, 2000). This indicates that the genesis and temperature of equilibration of the eclogite were older and thus unrelated to the Proterozoic orogeny which affected the granulites (Schmitz and Bowring, 2000). Therefore, the authors concluded that the projection of eclogite temperatures onto a P–T array defined by granulite thermobarometry (Pearson et al., 1995) resembling a hot oceanic geotherm resulting in P estimates

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of 1–2 GPa for the eclogitic samples may not be a suitable approach (Schmitz and Bowring, 2000). A cold cratonic geotherm seems to be more appropriate for the Archean eclogite and would yield much higher equilibration pressures of >3 GPa (Schmitz and Bowring, 2000). Furthermore, the eclogites show textural, petrographical and geochemical similarities to the pyroxenites of the xenolith suite suggesting a cogenetic origin representing subducted oceanic crust as indicated by oxygen isotope data (Tinguely et al., 2008). Since the reviewed studies on off-craton eclogites (Pearson et al., 1995; Schmitz and Bowring, 2000; Tinguely et al., 2008) are not based on identical samples (only the kimberlite localities are the same) the presence of crustal eclogites cannot be totally excluded. With this example we want to emphasize that the projection of xenoliths onto any type of geotherm can lead to erroneous results and that eclogitization of granulites at rather low P of 1–2 GPa rarely happens. We will discuss this conclusion below in the context of xenolith bulk compositions and expected equilibrium phase relations. The high temperatures of the xenoliths may also reflect a recent heat input (e.g. magmatism, regional metamorphism, shear heating) at the crust mantle boundary which has not reset most of the isotope ages. Higher temperatures may be generated by shear heating (Weinberg et al., 2007; Hartz and Podladchikov, 2008; Lu et al., 2011), hot fluids (Camacho et al., 2005) or melts at the crust mantle boundary zone and would also explain the high temperatures obtained for the spinel-bearing mantle xenoliths (see Section 2.3.). Furthermore, localized, transient heating may have only little effect on the surface heat flow thus preserving an overall cold geotherm. Geochronology of the studied areas suggests a time span of ≥1 Ga between the latest recorded metamorphic or magmatic event and the magmatic activity responsible for xenolith entrainment (Table 1; Fig. 2). Assuming that there were no unrecorded thermal events affecting the lower crust and upper mantle during this time span, the granulites were largely metastably preserved in the lower crust. While major element diffusion is supposed to occur on a time scale of several ten thousand years (Griffin et al., 1999), it has also been shown that purely thermally driven diffusion without the presence of a free fluid does not result in a complete isotopic homogenization (Glodny et al., 2008) and that the availability of fluid is a major factor limiting the extent of reactions and the attainment of equilibrium (e.g. Austrheim, 1987; Wayte et al., 1989; John and Schenk, 2003). The significant deviation of thermobarometric lower crustal xenolith data from a cratonic geotherm may thus be the result of a lack of equilibration due to insufficient temperatures, localized metasomatic alteration causing chemical disequilibrium or a relatively recent heat input and insufficient time for a re-equilibration to ambient P–T conditions. 3. Thermodynamic calculations and extraction of geophysical rock properties: comparison of granulitic and eclogitic compositions We have demonstrated that all crustal granulites from the studied areas yield P–T estimates which are not in agreement with cold cratonic geotherms (Fig. 1). While the processes causing the discrepancy between the thermobarometric results and the expected lower crustal conditions are still a matter of debate, it is possible to examine and quantify the effects of different P–T

Fig. 3. a) Phase diagram for a dry mafic granulite xenolith (37–40) from Kola with a composition given by Kempton et al. (2001) superimposed on a P–T-density diagram for the granulite. The box represents the P–T range of all Kola granulites based on geothermobarometry by Kempton et al. (1995). Small black circles mark individual P–T estimates for granulitic samples of similar composition as sample 37–40. White numbered circles mark the conditions of 550 °C at 1.2–1.4 GPa (numbers 1, 3 and 5; cold conditions expected along the cratonic geotherm) and 750 °C at 1.2–1.4 GPa (numbers 2, 4 and 6; range of P–T estimates of the granulite xenoliths) with the extracted densities and velocities listed in Table 2. b) P-wave velocity for the mafic granulite. c) Phase diagram of the Pl-poor granulite (N43). Composition listed in Kempton et al. (1995). d) VP for the Pl-poor granulite. e) Phase diagram based on a plagioclase and quartz-rich section (leucosome) of a migmatitic granulite (N57L) described by Kempton et al. (1995). f) VP for the leucosome. All mineral abbreviations according to Whitney and Evans (2010). Density and velocity data available as electronic supplement.

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conditions on geophysical properties of granulitic rocks comprising a majority of the lower crustal xenoliths as compared to eclogitic assemblages expected in the lower crust according to cratonic geothermal models. In the following section, we use whole rock compositions of lower crustal xenoliths for thermodynamic calculations to study the following: 1) Difference in densities and seismic velocities between granulitic and eclogitic assemblages (reported versus expected conditions) and consequences for interpretation of geophysical data; 2) Examination of the petrological feasibility of a proposed delamination model in a cratonic setting (Fennoscandian Shield) 3); The effect of fluids (metasomatism) on mineral assemblages and consequences for geophysical rock properties; 4) Influence of fluid-induced melting and subsequent melt extraction on rock densities; and 5) P–T-density relation of UHT granulites (Kaapvaal). 3.1. Suitability of available lower crustal xenoliths Xenolith-based studies are associated with a number of uncertainties. Xenoliths are random fragments which could have been located anywhere in the crustal column. Even if the samples showed complete textural equilibrium (which is often not the case) a significant uncertainty in geothermobarometry makes the determination of a derivation depth quite inaccurate. Furthermore, many exposed granulite terranes, including the ones in the proximity of the selected cratonic areas, show retrograde P–T paths implying decompression and/or cooling of the rocks after peak granulite conditions (e.g. Harley, 1989; Perchuk et al., 2000). Therefore, the rocks could have been moved within the crust after granulite-facies metamorphism and neither P nor T estimates recorded by the peak granulite assemblage are representative for the real position of the samples in the crust at the time of xenolith entrainment. However, numerous seismic wideangle velocity transects for the Fennoscandian Shield show that P-wave velocities do not exceed 6.70 km/s in the upper 25 km of the crust (e.g. Kuusisto et al., 2006). The majority of vP values calculated from modal compositions of mafic garnet granulites by Kuusisto et al. (2006), however, exceed velocities of 6.90 km/s. Furthermore, a lithological interpretation of crustal velocity profiles predicts the presence of mafic garnet granulites only below 25 km depth (Kuusisto et al., 2006). Combining seismic as well as petrological observations and geothermobarometric results, it is fairly reasonable to assume that significant amounts of mafic granulites are limited to below 25 km in the Karelian lower crust. Although the majority of amphibole-bearing granulites yield pressures of ~1.0 GPa we cannot exclude that some amphibolites have been derived from shallower levels since their lower seismic velocities do not provide the same constrictions as for the granulites. We are not aware of a comparable lithological interpretation study for the crust below Kola, but seismic studies show a crustal structure similar to the Karelian craton (e.g. Luosto, 1997). The density of the lowermost crust beneath the central, Archean part of Kaapvaal craton is estimated to 2860 kg/m 3 at seismic velocities in the range 6.4–6.7 km/s which excludes larger amounts of significantly denser rocks above the seismic Moho located at 35–40 km (Durrheim and Green, 1992; Nguuri et al., 2001; Niu and James, 2002). The Namaqualand lower crust yields velocities between 6.2 and 6.9 km/s also suggesting that the presence of mafic granulites is limited to the lower parts of the crust (Green and Durrheim, 1990; Durrheim and Green, 1992). All of our modeled phase diagrams are calculated over a range of pressures and temperatures based on the whole rock composition of the selected lower crustal xenoliths. They are therefore independent of the geobarometric results which have been presented in Fig. 1, Table 1 and Section 2.2. For all compositions, we have chosen a wide P–T range so that all relevant metamorphic facies are represented on the diagram. The actual P–T estimates for the relevant samples have only been used to verify that the model assemblage conforms to

the observed mineral assemblage. We use whole rock compositions of xenoliths obtained by previous workers. Where eclogites are not represented in the xenolith suite or where whole rock compositions for eclogites have not been provided, we have to rely on the compositions of a mafic granulite, the closest possible alternative. This approach is justifiable since: i) the lower crust is regarded to predominantly consist of mafic rocks (e.g. Rudnick and Fountain, 1995); ii) a mafic rock is the densest possible under the considered conditions; iii) we are able to cover the eclogite facies conditions with our modeled phase diagram and iv) we consider delamination due to phase changes in rocks of very similar (mafic) composition and not a densification due to compositional variation in the lower crust. Many models only use a single eclogite density (e.g. 3515 kg/m 3) ignoring the fact that rock densities are a function of P and T and therefore will show a significant variation even within the eclogite facies field of the same composition. Our model accounts for these changes in density and thus is a more suitable approach to discuss phase transitions, densification and possible delamination under varying conditions. 3.2. Methods Pseudosections as well as densities and seismic velocities have been calculated with the Gibbs free energy minimization software Perple_X 07 (Connolly and Kerrick, 1987; Connolly, 1990, 2009; Connolly and Petrini, 2002) using the thermodynamic dataset of Holland and Powell (1998 as revised by the authors in 2002). Solid solution models were chosen to match the rock lithologies as closely as possible within the given P–T estimates for each sample (other Perple_X details as well as the calculated data are available as electronic supplement). For reasons of simplification, MnO and P2O5 were excluded and all iron was assumed to be divalent. All dry pseudosections are calculated in the system TiO2–Na2O–CaO–K2O– FeO–MgO–Al2O3–SiO2 (TiNCKFMAS) except for the sapphirinebearing granulite where TiO2 was excluded. Where considered, fluid is represented by pure H2O. SiO2 saturation has been assumed for all rocks where a SiO2 phase (Qtz) has been reported to be part of the assemblage and is indicated in the top right corner of each phase diagram. A standard error in the order of 3% will be assumed for all computed densities and velocities as estimated by Connolly and Kerrick (2002). 3.3. Results 3.3.1. Kola The heterogeneous nature of the Kola xenoliths makes the samples ideal subjects to study differences in density and seismic velocity of granulitic and eclogitic assemblages for rocks showing compositional variations. Therefore, three xenoliths described by Kempton et al. (1995, 2001) have been chosen for the calculation of phase diagrams and the extraction of physical rock properties: sample 37–40 classified as a garnet granulite (Fig. 3a; b), sample N43 referred to as Pl-poor granulite (Fig. 3c; d) and sample N57L characterized as a feldspar-rich leucosome of a migmatitic garnet granulite (Fig. 3e; f). All samples are assumed to be essentially dry based on the description of the mineral assemblages by Kempton et al. (2001). They describe only accessory amounts of water-bearing minerals in these three samples, which do not have a significant influence on physical properties. Densities and seismic velocities (vP) are highlighted for six P–T conditions (Table 2; Fig. 3). The first three points with lower T roughly correspond to conditions expected along the cratonic geotherm while the last three represent P–T values in accordance with thermobarometric calculations of the xenoliths. The differences in density between “cold” and “hot” conditions are largest for the granulitic composition at elevated pressures (~ 360 kg/m 3) and

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Table 2 List of extracted densities and difference in densities at low and high temperatures and selected pressures calculated for the dry granulite, Pl-poor granulite and leucosome compositions given by Kempton et al. (1995, 2001). Numbers refer to the conditions represented by the white numbered circles in Fig. 3.

Mafic granulite 37–40; density Mafic granulite 37–40; P-velocity Pl-poor granulite N43; density Pl-poor granulite N43; P-velocity Leucosome N57L; density Leucosome N57L; P-velocity

550 °C, 1.2 GPa (1)

750 °C, 1.2 GPa (2)

Δρ/Δv

550 °C, 1.3 GPa (3)

750 °C, 1.3 GPa (4)

Δρ/Δv

550 °C, 1.4 GPa (5)

750 °C, 1.4 GPa (6)

Δρ/Δv

3040 kg/m3

2980 kg/m3

60 kg/m3

3330 kg/m3

2990 kg/m3

340 kg/m3

3380 kg/m3

3020 kg/m3

360 kg/m3

7.0 km/s

6.8 km/s

0.2 km/s

7.7 km/s

6.9 km/s

0.8 km/s

7.9 km/s

6.9 km/s

1.0 km/s

3650 kg/m

3

3550 kg/m

3

100 kg/m

3

3660 kg/m

3

3580 kg/m

3

80 kg/m

3

3660 kg/m

3

3610 kg/m

3

50 kg/m3

8.2 km/s

7.9 km/s

0.3 km/s

8.2 km/s

8.0 km/s

0.2 km/s

8.2 km/s

8.1 km/s

0.2 km/s

2880 kg/m3

2840 kg/m3

40 kg/m3

3020 kg/m3

2850 kg/m3

170 kg/m3

3140 kg/m3

2860 kg/m3

280 kg/m3

6.7 km/s

6.4 km/s

0.3 km/s

7.1 km/s

6.5 km/s

0.6 km/s

7.4 km/s

6.5 km/s

0.9 km/s

increase with rising pressure whereas they are significantly lower for the Pl-poor granulite at higher pressures (~ 50 kg/m 3) due to a decrease of Δρ with rising P (Table 2). The same trend is demonstrated by P-wave velocities with the largest change for the granulite at elevated pressures (Δv ~ 1.0 km/s; Table 2). 3.3.2. Karelia Due to the wide range of the modal abundances of water-bearing minerals (especially amphiboles), the Karelian xenoliths are regarded to be suitable to study the effects of water (simplified approach for a fluid) on geophysical rock properties. Furthermore, a model for eclogitization and subsequent delamination has been proposed for the Fennoscandian Shield (Kukkonen et al., 2008) making the Karelian samples the best available subjects to test the petrological feasibility of this model. In addition, the effect of dehydration melting and subsequent melt extraction on rock densities and P-wave velocities will be evaluated with help of the granulites from Karelia. The starting whole rock composition is provided by a mafic xenolith such as sample 7-HH-95 (Hölttä et al., 2000) since it contains only small amounts of amphibole and is therefore a good approximation to a water-free composition. Hölttä et al. (2000) calculated P–T conditions of ~ 840 °C and 1.0 GPa using mineral core compositions but we have taken somewhat higher pressures of 1.1 GPa due to the P correction suggested by Peltonen et al. (2006). At these reference

conditions, the calculated mineral assemblage has mineral modes similar to those given by Hölttä et al. (2000) and yields a rock density of 3230 kg/m 3 (Table 3; Fig. 4). The P-wave velocity of 7.20 km/s (Table 4) is reasonably close to vP of ~ 7.24 km/s calculated from modal compositions by Kuusisto et al. (2006). In order to study the densification of the mafic rock across the granulite to eclogite transition, we have extracted densities for pressures ranging from 1.4 to 2.0 GPa at 900 °C and 1000 °C (Table 3; Fig. 4). The dry mafic granulite reaches the eclogite-facies assemblage above 1.8 GPa at 900 °C (Table 3; Fig. 4a) but even at 2.0 GPa the rock yields a density below 3515 kg/m 3 (the reference eclogite density used by Kukkonen et al., 2008). Despite the lack of eclogite formation at most conditions considered, the equilibrated mafic composition can still yield densities equal or higher than the mantle density of 3320 kg/m 3 (indicated by bold values; Table 3). At higher T of 1000 °C, all calculated densities are lower at the equivalent pressures due to the effects of thermal expansion and therefore higher pressures are needed for the granulites to exceed the density of the mantle (Table 3). We have used the density values to calculate a depth estimate according to the formula: h = P / gρ (Table 3). The reference conditions of the mafic granulite correspond to a depth of ~ 40 km while the density of the mantle is exceeded at ~ 55 and 60 km for 900 and 1000 °C, respectively (Table 3). Since partial equilibration is very likely, we have also determined the proportion of

Table 3 Calculated densities for the Karelian mafic granulite (Fig. 4) with varying amount of water. Pressure to depth conversion with P–T dependent density and a constant density or 2700 kg/m3 and proportion of equilibrated rock needed to exceed the mantle density for metastable granulites with “higher” and “lower” densities. Bold numbers indicate the first density equal to or exceeding the mantle density of 3320 kg/m3. Numbers correspond to white circles in Fig. 4.

840 °C, 1.1 GPa (1) 900 °C, 1.4 GPa (2) 900 °C, 1.5 GPa (3) 900 °C, 1.6 GPa (4) 900 °C, 1.7 GPa (5) 900 °C, 1.8 GPa (6) 900 °C, 1.9 GPa (7) 900 °C, 2.0 GPa (8) 1000 °C, 1.4 GPa 1000 °C, 1.5 GPa 1000 °C, 1.6 GPa 1000 °C, 1.7 GPa 1000 °C, 1.8 GPa 1000 °C, 1.9 GPa 1000 °C, 2.0 GPa 600 °C, 1.5 GPa (9) a b

7-HH-95 Density [kg/m3]

Depth h = P/g∗ ρ [km]a

% to exceed 3320 kg/m3; low ρ of granulite

% to exceed 3320 kg/m3; high ρ of granulite

7-HH-95, 0.2 wt.% H2O density [kg/m3]

7-HH-95, 0.5 wt.% H2O density [kg/m3]

3230 3270 3290 3320 3370 3430 3460 3470 3250 3260 3290 3310 3340 3390 3450 3490

37 48 51 54 58 61 65 68 48 51 54 58 61 65 68 51

– – – 100 (3150)b 80 (3150) 60 (3160) 50 (3160) 50 (3160) – – – – 90 (3150) 70 (3150) 60 (3150) –

– – – 100 (3270) 50 (3270) 30 (3280) 20 (3280) 20 (3290) – – – – 70 (3270) 40 (3270) 30 (3280) –

3190 3250 3280 3310 3350 3410 3450 3450 3230 3250 3270 3290 3330 3380 3430 3490

3150 3230 3250 3280 3330 3390 3420 3430 3210 3230 3250 3270 3300 3350 3410 3460

g = 9.81 m/s2; ρ = 3000 kg/m3. The density of the granulite used for each calculation is given in brackets and is the result of the P–T dependent properties of the specified minerals (Hacker and Abers, 2004).

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dense granulite/eclogite in the lower crust which is needed for the rocks to exceed mantle density (Table 3). At the selected P–T conditions, we use the densities calculated for the mafic rock assuming complete equilibration (Table 3; Fig. 4) and mix them with a high and low density value of a preserved metastable granulite until the average mantle density of 3320 kg/m 3 is exceeded. The result is highly variable (Table 3) depending on densities used for the granulite and the P–T conditions. To evaluate the effect of water on reactions, densities and delamination we consider the addition of 0.2 and 0.5 wt.% water to the initially dry composition of sample 7-HH-95 which predicts the formation of ~ 10 and ~ 18–20 vol.% of amphiboles, respectively, in the P–T stability field of the Karelian samples (Fig. 4b and c). The presence of amphibole will reduce the rock density by ~ 40 kg/m 3 (0.2 wt.% water) and ~ 80 kg/m 3 (0.5 wt.% water) at the reference conditions of 840 °C and 1.1 GPa (Table 3; Fig. 4). However, at elevated temperatures (800–900 °C) water-bearing phases are not part of a stable mineral assemblage resulting in a dry granulitic composition in addition to free water (Fig. 4b and c). If we exclude melting momentarily and assume very high rock permeabilities, the water can leave the system completely and the resulting rock densities will be identical to the ones of the dry granulite (7-HH-95; Table 3). If the rocks are highly impermeable, however, some or all fluid may remain in the pores of the rock (e.g. Simon, 2011) and we will have to include the fluid in the computation of aggregate properties. This will reduce the density of the system by ~ 0.5% from a dry granulite to a rock containing 0.2 wt.% water and by ~ 1.2% from dry to 0.5 wt.% water (Table 3; Fig. 4b and c). The density at eclogite-facies conditions (Table 3; Fig. 4) is slightly reduced by ~ 30 kg/m 3 from the dry granulite to the composition containing 0.5 wt.% water due to the formation of phengite. To study the effect of partial melting, we have calculated the amount of melt forming in a mafic granulite with 0.2 and 0.5 wt.% water, respectively (Fig. 5). As expected, the stability field and the amount of melt will increase with increasing water content in the bulk composition (Fig. 5). Most melt is generated at the lowest pressure and 1000 °C while the lowest amounts of melt are present at the highest pressure and 900 °C (Fig. 5). Since the first melt is generally expected to be more felsic, the density of the restite is expected to increase if this melt is drained from the residue. To quantify this effect, we assume the rocks to first reach a certain pressure (starting at 1.4 GPa up to 2.0 GPa), then allow the melt to form and leave the system along this isobar and finally extract the densities at 900 and 1000 °C. The resulting increase in density from the dry granulite to the composition containing 0.2 wt.% water is rather insignificant with a maximum densification of ~ 0.5% (Table 4). With a higher water content of 0.5 wt.%, this value increases to ~ 1%. If the water content is increased to 2 wt.% (water-saturation) the densification is up to 3% due to the melt extraction (Table 4). For compositions with small amounts of water, the densification is most distinctive at high pressures and temperatures (1000 °C, 2.0 GPa). This trend is reversed for the composition with 2.0 wt.% water with the highest increase in density relative to the dry granulite at 1.4 GPa and 900 °C. Fig. 4. a) Phase diagram for a dry mafic granulite xenolith from Karelia with a bulk composition as given by Hölttä et al. (2000) superimposed on the P–T-density diagram. The box represents the P–T range of all Karelian granulites based on geothermobarometry by Hölttä et al. (2000). The black circle marks the P–T estimates for this sample. White circles represent conditions at which densities (Table 3) have been extracted. 1: 840 °C and 1.1 GPa; 2–8: 1.4–2.0 GPa at 900 °C; 9: 600 °C, 1.5 GPa. The arrow indicates retrogression of eclogite to granulite when the temperature is increased. b) Phase and density diagram for the mafic granulite composition of Fig. 4a with 0.2 wt.% water without using a melt solid solution. Extremely small phase fields have been omitted for reasons of clarity since they are not of significant importance for the extracted rock properties. Circles mark P–T conditions of samples with ~10 vol.% amphiboles (Hölttä et al., 2000) c) Phase and density diagram for the mafic granulite with 0.5 wt.% water. Circles mark samples with ~20 vol.% amphiboles (Hölttä et al., 2000).

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Table 4 Calculated densities for the mafic granulite from Karelia with varying amounts of water (Fig. 4). Densities for the compositions with 0.2, 0.5 and 2.0 wt.% are calculated after melt extraction. Densification is given relative to the density of the dry granulite. Bold numbers indicate the first density equal to or exceeding the mantle density of 3320 kg/m3. 7-HH-95; density 840 °C, 1.1 GPa (1) 900 °C, 1.4 GPa (2) 900 °C, 1.5 GPa (3) 900 °C, 1.6 GPa (4) 900 °C, 1.7 GPa (5) 900 °C, 1.8 GPa (6) 900 °C, 1.9 GPa (7) 900 °C, 2.0 GPa (8) 1000 °C, 1.4 GPa 1000 °C, 1.5 GPa 1000 °C, 1.6 GPa 1000 °C, 1.7 GPa 1000 °C, 1.8 GPa 1000 °C, 1.9 GPa 1000 °C, 2.0 GPa

3230 3270 3290 3320 3370 3430 3460 3470 3250 3260 3290 3310 3340 3390 3450

kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3

7-HH-95, 0.2 wt.% H2O; density 3250 3280 3300 3330 3380 3440 3480 3480 3260 3270 3300 3320 3350 3400 3470

kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3

Densification ~%

7-HH-95, 0.5 wt.% H2O; density

0.4 0.3 0.3 0.3 0.3 0.4 0.4 0.4 0.3 0.3 0.3 0.3 0.3 0.4 0.5

3260 3300 3320 3350 3390 3460 3490 3500 3270 3290 3310 3330 3370 3420 3480

Thus, depending on the amount of water in the rock, melting and subsequent melt extraction can lead to a significant densification of the residual crust. 3.3.3. Kaapvaal Due to their exceptional mineralogy and P–T estimates, we have chosen an UHT sapphirine-bearing granulite for a density calculation. The composition has been calculated by Dawson and Smith (1987) from mineral modes and phase analyses. The authors pointed out that the mineral modes are approximate and not necessarily representative due to the small size of the xenoliths (Dawson and Smith, 1987) which has some consequences on the calculated density diagram. Nevertheless, the chosen sample (specimen E) yields exceptional high densities (Fig. 6) which will not be altered significantly by estimation errors. Within the P–T estimates (>1040 °C, 0.9–1.1. GPa; Dawson et al., 1997), we have tried to find three conditions where the mineral modes are as close as possible to the original estimates (Table 5). The two densest minerals (garnet and sapphirine) coincide within ± 4 vol.% while sillimanite is the only phase which does not comply with the modal estimate (Table 5). Picking a lower value, however, would result in lower amounts of garnet as well which has a greater effect on the rock density. The density varies little for the three conditions and is with ~ 3340–3350 kg/m 3 (Table 5) closer to mantle densities than expected crustal density values of 2860 kg/m 3 in the Kaapvaal lower crust (Niu and James, 2002). 4. Discussion 4.1. Consequences of metastability for interpretation of rock types The assumption that cratonic lower crustal rocks have equilibrated to eclogite-facies conditions while the data implies that metastable granulites predominate (Section 2) may result in the misinterpretation of rock types if only geophysical data are available. The density reduction between the eclogitic assemblage of the mafic rock from Kola (550 °C; 1.4 GPa; Table 2; Fig. 3a) and the granulite-facies equivalent for the same composition (750 °C; 1.4 GPa; Table 2; Fig. 3a) amounts to ~ 11%. In fact, a rock of more felsic composition (e.g. a dry pelite such as mix1p in Semprich et al., 2010) equilibrated to 550 °C and 1.3 GPa will have almost the same density as the mafic granulite (~ 2990 kg/m 3; Table 2) equilibrated at 750 °C and 1.3 GPa. P-wave velocities for both compositions at the same conditions show a similar trend with the metastable granulitic rock even yielding a slightly lower value (~ 6.9 km/s) than the dry metapelitic rock equilibrated to eclogite facies conditions (~ 7 km/s). A remarkable example of metastably preserved rocks in the lower crust is the UHT sapphirine-bearing granulites. In addition to

kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3

Densification ~%

7-HH-95, 2.0 wt.% H2O; density

Densification ~%

1.0 0.8 0.8 0.8 0.8 0.9 0.8 0.9 0.8 0.8 0.7 0.7 0.8 0.8 1.0

– 3370 3390 3410 3450 3500 3510 3520 3340 3360 3360 3380 3410 3470 3500

– 3.0 2.9 2.5 2.4 2.2 1.4 1.5 2.9 2.8 2.4 2.2 2.1 2.2 1.6

kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3 kg/m3

recording extremely high temperatures, they also yield densities highly exceeding the estimate for the lower crust below Kaapvaal and are rather similar to mantle densities. Since they do not seem to significantly disturb seismic or gravity data, however, these rocks might be scarce. Alternatively, they may represent the strong reflector in some areas which is commonly interpreted as the Moho. The issue of metastability therefore adds an additional complication to the already existing ambiguity of deriving rock types from geophysical data. 4.2. Effect of P, T and composition on densities The simplified approach of taking a constant density value for each rock type at all conditions is not sufficient to adequately explain the complex processes in the lower crust. Density is strongly dependent on pressure, temperature and composition (e.g. Semprich et al., 2010) and we have demonstrated that eclogites forming at higher temperatures yield lower densities (Table 3; Fig. 4) than the frequently used value of 3515 kg/m3 (Christensen and Mooney, 1995). On the other hand, some non-eclogitic compositions can reach a density above 3500 kg/m 3 such as the Pl-poor granulite from Kola (Table 2, Fig. 3b). 4.3. Consequences for gravitational stability/delamination The very small or completely lacking density contrast at the Moho as well as the presence of garnet and clinopyroxene xenocrysts with compositions typical for crustal eclogites is seen as proof for the presence of eclogites and delamination of parts of the Karelian lower crust (Kukkonen et al., 2008). However, this process requires a sufficiently low viscosity of the upper mantle and therefore high temperatures (900–1000 °C) as well as a sufficient positive density contrast between the lowermost crust and the upper mantle (Kukkonen et al., 2008). In the case of incomplete eclogitization, the authors assume the lower crust to consist of a mixture of mafic granulite (3170 kg/m 3) and mafic eclogite (3515 kg/m 3) and conclude that the mafic layer would still sink into the mantle if the eclogite proportion exceeds 44% (Kukkonen et al., 2008). Kukkonen et al. (2008) propose two scenarios: 1) initially hot lower crust and upper mantle allowing for rapid eclogitization and delamination immediately after crustal thickening; and 2) initially low temperatures in the crust and upper mantle and subsequent slow heating of the stacking plates. Our calculations demonstrate that a number of petrological restrictions add complications to the feasibility of delamination at relatively high temperatures: 1) Eclogitization is shifted to higher pressures at elevated temperatures. At moderately low temperatures (500–600 °C), eclogites can form at 1.5 GPa (assuming complete

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Fig. 6. Phase and density diagram for a UHT sapphirine granulite. Black circles represent condition listed in Table 5.

and 2.0 GPa (~ 70 km; Table 3); 4) Retrogression of eclogite to granulite during heating of the crust. If eclogites are present in the lower crust which could not delaminate due to unfavorable mantle viscosities, it is very likely that they will at least partially react back to granulites since element diffusion is more efficient with increasing temperature (as demonstrated by the arrow in Fig. 4a); and 5) Presence of less dense rock types. In summary, a sufficient densification of the lower crust for delamination at relatively high temperatures is unlikely above a depth of 60 km if pressures are only lithostatic. The lacking density contrast between the lower crust and upper mantle could also be a result of magmatic underplating and the crystallization of granulites with densities ~ 3320 kg/m 3 at a depth of ~ 55 km (Table 3). The presence of xenocrysts is not necessarily a proof for the existence of eclogites as shown by a study of the deep crust beneath the Dharwar craton (India) where garnet and clinopyroxene in mafic granulite xenoliths show compositional similarity to crustal eclogites (Dessai et al., 2010). 4.4. Influence of water on reactions/phase petrology

Fig. 5. a) Phase diagram and melt proportion in vol.% for the Karelian granulite with 0.2 wt.% water. Labels of extremely small fields have been omitted. Circles indicate the amount of melt at 840 °C at 1.1 GPa, 900 °C at 1.4 and 2.0 GPa and 1000 °C at 1.4 and 2.0 GPa. b) Phase diagram and melt proportion for the granulite with 0.5 wt.% water.

equilibration) which is equivalent to ~ 50 km depth (Table 3; Fig. 4a). The higher temperatures (900–1000 °C) required for a sufficiently low viscosity of the mantle will shift the formation of eclogites to depths of ~ 65–70 km (P: 1.9–2.0 GPa; Table 3); 2) Rocks equilibrated at higher temperatures yield lower densities than at lower T. Even when transformed to eclogite the mafic composition will not reach densities above 3500 kg/m 3 within the applied P–T conditions (Table 3). Pressures above 2.0 GPa (depths > 70 km) are required until comparable densities are reached; 3) Incomplete reactions and metastability will delay the densification towards higher pressures/ greater depth. While the density of the equilibrated rock is defined by the P–T conditions, the metastable rock can display a range of densities depending on the conditions at which it stopped equilibrating. Assuming a density close to the mafic granulite xenolith, the proportion of equilibrated rocks has to amount to 50–60% at 900–1000 °C

Aqueous (water-bearing) fluids can change the rock chemistry by transporting elements to or from the rock, induce the formation of hydrous minerals which will affect its petrophysical properties, catalyze sluggish mineral reactions when introduced to a metastably preserved rock and reduce the melting temperature. It is beyond question that not all rocks are exposed to the same amounts of fluids. But can highly varying proportions of amphibole from almost zero to ~ 60 vol.% as in the Karelian xenoliths (Hölttä et al., 2000) be a result

Table 5 Modal estimates of the UHT sapphirine granulite specimen E (Dawson and Smith, 1987) and calculated mineral modes at three conditions. Numbers correspond to black circles in Fig. 6.

Spr Grt Qtz Pl Sil Kfs ρ [kg/m3]

E

1070 °C; 0.95 GPa (1)

1070 °C; 0.96 GPa (2)

1080 °C; 0.97 GPa (3)

40 30 15 5 5 Present –

38 28 16 5 13 0.2 3340

37 28 16 5 14 0.2 3350

37 28 16 5 14 0.2 3350

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of varying water content alone? We could reproduce the modal amounts of amphiboles for a number of xenoliths by varying the amount of water added to the mafic composition (Fig. 4b and c) but even a water-saturated composition is predicted to contain only ~ 26 vol.% amphiboles. We cannot exclude the possibility that the whole-rock compositions of the xenoliths are not representative due to their relatively small size. Alternatively, the higher modal amounts of hornblende may be the result of the granulites being exposed to changed P–T conditions (since the geothermobarometric results have only been obtained for the dry granulitic assemblage), more than one fluid infiltration and/or the influence of potassiumbearing amphibole which could not be modeled with Perple_X due to the lack of an adequate solid solution model. The formation of water-bearing minerals and reduction of densities are limited to the low-medium T range. Once temperatures over ~ 750 °C (for the studied rock compositions) are reached, the fluid tends to be present as a free phase facilitating dehydration melting. Adding water to a rock at high temperatures will therefore primarily cause melt formation and only trigger eclogitization at sufficiently high pressures (~ 1.9 GPa). 4.5. Comparison with recent concepts of craton formation and evolution Models for craton formation have to account for cold, refractory, buoyant Archean lithospheric mantle on the one hand and broadscale crustal melting responsible for the post-orogenic granitic magmatism on the other. Furthermore, coinciding ages from mantle and lower crustal xenoliths as well as surface rocks suggest long-time coupling of the Archean crust and underlying mantle lithosphere, in particular for the Fennoscandian Shield and the Kaapvaal craton (e.g. Richardson et al., 1984; Menzies et al., 1999; Griffin et al., 2004; Peltonen et al., 2006). Moreover, many cratons are characterized by seismic refraction velocities in the range of 6.8–7.0 km/s which is interpreted to correspond to intermediate bulk compositions hence excluding the presence of a mafic layer with high velocities in the lower crust (e.g. Durrheim and Mooney, 1994; Rudnick and Gao, 2003). The refractory and buoyant nature of the mantle results mainly from the lack of garnet in lithospheric mantle peridotite and is commonly attributed to high degrees of melt extraction due to mantle plumes or at plate boundaries such as mid-oceanic ridges and subduction settings (e.g. Herzberg, 1993; Canil, 2002; Carlson et al., 2005). However, a study of inclusions in diamonds from the lithospheric mantle of several cratons reports only peridotitic material prior to 3.2 Ga while diamonds from younger subcontinental mantle also contain an eclogitic component (Shirey and Richardson, 2011). The sudden appearance of eclogitic mineral inclusions at ~ 3 Ga has been interpreted as the onset of plate tectonics which would exclude the formation of older Archean lithosphere in a mid-ocean ridge or subduction environment but instead support the creation of continental crust over either upwelling or downwelling mantle (Shirey and Richardson, 2011; Van Kranendonk, 2011). Nevertheless, significant reworking of the lithospheric mantle and the formation of the exceptional thick mantle keel could have happened by means of tectonic processes after ~ 3 Ga (e.g. Schmitz et al., 2004). Among the processes which have been proposed for the widespread late granitic magmatism in many cratons (thermal relaxation after orogenic thickening, e.g. England and Thompson, 1986; radiogenic heating, e.g. Davies, 1995; delamination of mantle lithosphere, e.g. Smithies and Champion, 1999; underplating of basaltic magmas, e.g. Petford and Gallagher, 2001) only magmatic underplating and delamination of lithospheric mantle can explain the generation of substantial volumes of melt. However, instability of the mantle lithosphere and asthenospheric heating are inconsistent with the long-term buoyancy of the Archean mantle as well as the continuing coupling of the crust and mantle since the Archean (e.g. Percival and

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Pysklywec, 2007). Underplating of magma is commonly associated with the formation of a residual mafic layer in the lower crust characterized by high velocities and hence could explain the evolution of cratons where this feature is present (e.g. Karelia, Peltonen et al., 2006; North China, Zhang et al., 2012). Episodic magmatic activity is also corroborated by the diversity of zircon U–Pb ages in the same granulite xenoliths (e.g. Hölttä et al., 2000; Zhang et al., 2012), the presence of pyroxenites and banded (migmatitic) granulites (e.g. Kempton et al., 2001; Zhang et al., 2012) and the considerably modified mantle lithosphere (e.g. Peltonen and Brügmann, 2006; Zhang et al., 2012). The thick subcontinental mantle keel may then have formed as a result of continent collision and the emplacement of plate-like mantle lithosphere (underplating) with the rheologically stronger mafic crust remaining coupled to the underlying mantle (Gray and Pysklywec, 2010). For cratons lacking a high-velocity (mafic) layer in the lower crust, Percival and Pysklywec (2007) have proposed a late-orogenic, density-driven lithosphere inversion to account for the extensive granitic magmatism. The model requires a thick eclogitic layer (~20 km) in the lower crust with an assumed fixed density of 3500 kg/m 3. However, using the density diagram of the mafic granulite from Karelia (7-HH-95) and the parameters provided by Percival and Pysklywec (2007), e.g. a mafic layer from 40 to 60 km depth corresponding to a (lithostatic) pressure of ~ 1.2 to 1.8 GPa and 730 °C at the crust–mantle boundary, we have estimated a mean density of ~ 3390 kg/m 3 for the mafic layer. Despite assuming a homogeneous mafic layer and complete equilibration, the mafic composition will not be transformed to eclogite at P–T conditions corresponding to depths ~ 40 km unless the pressures are much higher than lithostatic (e.g. dynamic pressure due to collision). 5. Conclusions A majority of thermobarometric P–T estimates for cratonic mantle xenoliths correlates well with a cold conductive geotherm. In contrast, P–T estimates for lower crustal xenoliths deviate significantly from the cratonic geotherms. Independent from the individual cratonic setting, the temperatures are ~ 200–300 °C higher than what is expected at the base of the cratonic lower crust (~500–600 °C at ~ 1.3–1.6 GPa). Possible explanations for this deviation may be a lack of equilibration due to thermal closure and/or the lack of fluid. Alternatively, the high temperatures may be the result of a relatively recent localized heat input (e.g. shear heating, hot fluids or melts) which has not reset the rocks as indicated by the old isotope ages. A mafic granulitic assemblage which has been preserved metastably under eclogite-facies conditions has densities and velocities similar to a felsic composition equilibrated to eclogite-facies conditions (~ 2990 kg/m 3; ~ 7 km/s). Metastable UHT granulites with an uncommon composition can be as dense as the mantle (~3350 kg/m 3). Density is always a function of pressure, temperature and composition. Using a constant density value for any rock type will lead to inaccurate results. While high temperatures of 900–1000 °C are required to ensure suitable mantle viscosities for the delamination of crustal rocks, most rock compositions are not transformed to eclogite at pressures expected at the base of the lower crust (50–60 km). Slow reaction kinetics may preserve the granulite. Fluids can only trigger the eclogitization if the P–T conditions and the bulk rock composition allow for an eclogitic assemblage. Fluid-induced melting is more probable at elevated temperatures above 800 °C. Dehydration melting may cause a densification of the rocks of up to 3% depending on the degree of water saturation and assuming that the melt is efficiently removed. Eclogitization at lower temperatures (500–600 °C) is unlikely due to metastability while the formation of eclogite at higher T (>700 °C) requires pressures exceeding lithostatic in the lower crust. Hence, geodynamic models requiring a significant amount of eclogitization

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