the origin of shoshonites: new insights from the tertiary high-potassium intrusions of eastern tibet

ORIGINAL PAPER

The origin of shoshonites: new insights from the Tertiaryhigh-potassium intrusions of eastern Tibet

Ian H. Campbell • Aleksandr S. Stepanov •

Hua-Ying Liang • Charlotte M. Allen •

Marc D. Norman • Yu-Qiang Zhang • Ying-Wen Xie

Received: 6 February 2013 / Accepted: 25 October 2013

� Springer-Verlag Berlin Heidelberg 2014

Abstract The shoshonitic intrusions of eastern Tibet, which

range in age from 33 to 41 Ma and in composition from

ultramafic (SiO2 = 42 %) to felsic (SiO2 = 74 %), were

produced during the collision of India with Eurasia. The mafic

and ultramafic members of the suite are characterized by

phenocrysts of phlogopite, olivine and clinopyroxene, low

SiO2, high MgO and Mg/Fe ratios, and olivine forsterite con-

tents of Fo87 to Fo93, indicative of equilibrium with mantle

olivine and orthopyroxene. Direct melting of the mantle, on the

other hand, could not have produced the felsic members. They

have a phenocryst assemblage of plagioclase, amphibole and

quartz, high SiO2 and low MgO, with Mg/Fe ratios well below

the values expected for a melt in equilibrium with the mantle.

Furthermore, the lack of decrease in Cr with increasing SiO2

and decreasing MgO from ultramafic to felsic rocks precludes

the possibility that the felsic members were derived by

fractional crystallization from the mafic members. Similarly,

magma mixing, crustal contamination and crystal accumula-

tion can be excluded as important processes. Yet all members

of the suite share similar incompatible element and radiogenic

isotope ratios, which suggests a common origin and source. We

propose that melting for all members of the shoshonite suite

was initiated in continental crust that was thrust into the upper

mantle at various points along the transpressional Red River-

Ailao Shan-Batang-Lijiang fault system. The melt formed by

high-degree, fluid-absent melting reactions at high-T and high-

P and at the expense of biotite and phengite. The melts acquired

their high concentrations of incompatible elements as a con-

sequence of the complete dissolution of pre-existing accessory

minerals. The melts produced were quartz-saturated and reac-

ted with the overlying mantle to produce garnet and pyroxene

during their ascent. The felsic magmas reacted little with the

adjacent mantle and preserved the essential features of their

original chemistry, including their high SiO2, low Ni, Cr and

MgO contents, and low Mg/Fe ratio, whereas the mafic and

ultramafic magmas are the result of extensive reaction with the

mantle. Although the mafic magmas preserved the incompat-

ible element and radiogenic isotope ratios of their crustal

source, buffering by olivine and orthopyroxene extensively

modified their MgO, Ni, Cr, SiO2 contents and Mg/Fe ratio to

values dictated by equilibrium with the mantle.

Keywords Shoshonite � High-potassium magmas � High-

potassium intrusions � Eastern Tibet � Red River-Ailao

Shan-Batang-Lijiang fault system

Introduction

The origin of shoshonitic suites is controversial (Pecce-

rillo 1992). There is a general consensus that they were

Communicated by M. W. Schmidt.

Electronic supplementary material The online version of thisarticle (doi:10.1007/s00410-014-0983-9) contains supplementarymaterial, which is available to authorized users.

I. H. Campbell (&) � A. S. Stepanov � H.-Y. Liang �C. M. Allen � M. D. Norman

Research School of Earth Sciences, Australian National

University, Canberra, ACT 0200, Australia

e-mail: [email protected]

H.-Y. Liang

Key Laboratory of Mineralogy and Metallogeny, Guangzhou

Institute of Geochemistry, Chinese Academy of Sciences,

Guangzhou 510640, China

Y.-Q. Zhang � Y.-W. Xie

Laboratory of Marginal Sea Geology, Guangzhou Institute of

Geochemistry, South China Sea Institute of Oceanology,

Chinese Academy of Sciences, Guangzhou 510640, China

123

Contrib Mineral Petrol (2014) 167:983

DOI 10.1007/s00410-014-0983-9

http://dx.doi.org/10.1007/s00410-014-0983-9

derived from partial melting of mantle that had been

metasomatically modified by subduction-related fluids

(Morrison 1980; Edgar and Arima 1985; Foley and Pec-

cerillo 1992; Aoki et al. 1981; Jiang et al. 2002; Pecce-

rillo 1990; Nelson 1992; Turner et al. 1996). The high

concentrations of large ion lithophile elements (LILE),

light rare earth elements (LREE) and volatiles, which

characterize shoshonitic suites, are assumed to have been

introduced into their mantle source region by fluids

derived from the underlying subducted slab. Meen (1987)

suggested that shoshonites form by low degrees of melt-

ing of hydrous upper mantle lherzolite that had been

metasomatically enriched in LILE and LREE. The

metasomatic hypothesis for the origin of shoshonites is

consistent with the experimental work of Wyllie and

Sekine (1982) who showed that potassic magmas can be

produced by melting hydrated mantle. On the other hand,

post-collisional shoshonitic granitoids from the East

African Orogen (Kuster and Harms 1998) and the Kunlun

Orogenic Belt, Xinjiang (Jiang et al. 2002), are quartz-

saturated and therefore cannot be derived directly from

the mantle. It has been suggested that they were derived

by partial melting of the lower crust, following under-

plating by mantle-derived magma (Jiang et al. 2002;

Kuster and Harms 1998).

We describe a shoshonite suite of ultramafic through

basaltic to granitic rocks in eastern Tibet that follow the

Red River fault zone and its northern extension, the Ba-

tang-Lijiang fault system, a total strike length of over

1,500 km. The mafic and ultramafic members have crys-

tallized from magmas in equilibrium with the mantle,

whereas the felsic members are quartz-saturated and

therefore cannot be in equilibrium with mantle olivine.

Yet all of the intrusions from this suite share the same

‘‘crustal’’ radiogenic isotope and trace element signatures,

which suggest that all members of the suite share a similar

source and have a related origin. We will argue that the

common source was continental crust that was pushed into

mantle during the collision of India with Eurasia, where it

underwent partial melting. The resulting silica-saturated

melt was out of equilibrium with the overlying mantle and

reacted with it to produce garnet and orthopyroxene during

its ascent. We suggest that, in case of the shoshonitic

granitoids, the reaction was minimal so that the ascending

melts retained the essential features of their crustal source

region. In contrast, the ultramafic members of the suite

were produced by extensive reaction with the overlying

mantle so that they reached equilibrium with olivine and

orthopyroxene. As a consequence, the mantle controlled

their Ni, Cr and major element contents, whereas their

incompatible trace elements and radiogenic isotope char-

acteristics are largely derived from their initial crustal

source.

Geological setting

The Red River-Ailao Shan-Batang-Lijiang fault system in

eastern Tibet and western Yunnan forms part of a series of

north to north-west striking shears bounded to the east by the

Xiaojiang fault, to the south by the Red River shear zones and

to the west by the Gaoligong fault (Fig. 1). The movement

within this region was sinistral during the Tertiary and was

coeval with transpressive thrusting and folding in the red bed

basin of Yunnan. This deformation is thought to have started

prior to 42 Ma (Wang and Burchfiel 1997) and is therefore

coeval with the collision of India with Eurasia. The faults are

the accommodating structures that allowed India to move

north by pushing Southeast Asia to the south-east (Peltzer and

Tapponnier 1988; Tapponnier et al. 1982, 1986). Total

movement on the Red River-Ailao Shan fault is

700 ± 200 km (Leloup et al. 1995; Lacassin et al. 1996),

which is greater than the displacement on major plate

boundary faults such as the San Andreas and Alpine faults

(Leloup et al. 1995; Tapponnier et al. 2001). It deeps steeply at

60–70� towards north-east (Tapponnier et al. 1990) and

reaches the base of the lithosphere (Tapponnier et al. 2001).

The Moho depth is estimated at about 40–42 km on the

northeast side of the fault zone and about 30–37 km on the

southwest side (Xu et al. 2006), so the crust on the northern

side on the fault zone extends about 5 to 10 km into the

adjacent mantle.

The Red River-Ailao Shan-Batang-Lijiang fault system

is associated with the Ailao Shan-Red River metamorphic

belt, which has a sharp horizontal metamorphic gradient on

either side of its gneisses core. Unmetamorphosed rocks

crop out only 3–10 km away from the high-grade gneisses.

These steep thermal gradients could be due to frictional

heating within the shear zone or differential uplift between

the gneisses and the surrounding sedimentary rocks (Le-

loup et al. 1993, Leloup and Kienast 1993, Scharer et al.

1990, Scharer et al. 1994; Harrison et al. 1992, 1995).

The Red River shoshonitic series is a suite of 31–41 Ma

high-potassium igneous rocks that are localized along a

narrow 1,500-km-long belt that follows the Nanqian thrust

belt, the Batang-Lijiang fault system and the Red River-

Ailao Shan fault (Pan et al. 1990; Wang et al. 2001) in

eastern Tibet (Fig. 1). They consist of small pipes of

phenocryst-rich rocks that range in composition from

ultramafic to felsic, accompanied by minor volcanic rocks.

The intrusions were emplaced into Cenozoic siltstones and

conglomerates, Triassic–Jurassic limestones and clastic

rocks, Middle Silurian limestones and Carboniferous–

Permian limestones. They have chilled margin and meta-

morphic aureoles that consist of hornfels where the con-

tacts are with clastic sediments, and skarns where granitic

rocks intrude limestone. Most contacts are near vertical.

The ultrabasic and basic rocks occur as pipes with

983 of 22 Contrib Mineral Petrol (2014) 167:983

123

Fig. 1 Simplified tectonic map of the western Yunnan area of eastern

Tibet, (modified after Wang et al. 2001). Shown are the locations of

samples used in this and in previous studies, together with the age of

the intrusions in Ma. The number in round brackets, on both the map

and below, identify the samples, the square brackets give the dating

method and source of the data, sample numbers from Liang et al.

(2007) have the form Y81-974, Y83-740, etc. (1) [U–Pb, LS]; (2)

Y81-78 [Zircon, U–Pb]*; (3) [Ar, JH]; (4) [Ar, JH]; (5) Y83-803

[Zircon, U–Pb]*; (6) n.d.; (7) Y81-974 [Zircon, U–Pb]*; (8) Y83-634

Y83-620 [Zircon, U–Pb]*; (9) [K–Ar, Z]; (10) n.d.; (11) [K–Ar, Z];

(12) n.d.; (13) [K–Ar, Z]; (14) n.d.; (15) [K–Ar, Z]; (16) [K–Ar, Z];

(17) Y83-745 [Zircon, U–Pb]*; (18) Y83-740 [Zircon, U–Pb)*]; (19)

Y83-760 [Zircon, U–Pb]*; (20) Y83-760 [Zircon, U–Pb]*; (21) Y83-

791 [Zircon, U–Pb]*; (22) [U–Pb,SU]; (23) Y83-909 [Zircon, U–

Pb]*; (24) n.d.; (25) [Ar, JH]; (26) T83-85 [Zircon, U–Pb]*; (27) T83-

110 [Zircon, U–Pb]*; (28) T83-931 [Zircon, U–Pb]*; (29) T83-305

[Zircon, U–Pb]*; (30) T83-358 [Zircon, U–Pb]*; (31) T83-404

[Zircon, U–Pb]*; (32)T82-141 [Zircon, U–Pb]*; (33) T83-258

[Zircon, U–Pb]*; (34) n.d. Data sources LS, Zhang and Scharer

(1999); JH, Wang et al. (2001); Z, Zhang and Xie (1997); SU, Scharer

et al. (1990) [47]; * Liang et al. (2007) U–Pb zircon dates

Contrib Mineral Petrol (2014) 167:983 of 22 983

123

Table 1 Representative analyses of shoshonitic rocks from eastern Tibet

Sample 86–68 86–132 86–8 86–190 86–002 86–80 86–101 81–909 86–40Locality 11 13 12 16 15 10 13 23 6

SiO2 43.80 44.82 50.17 51.02 52.26 52.32 55.04 57.81 58.24

TiO2 0.74 0.88 0.61 0.58 0.64 0.66 0.61 0.70 0.47

Al2O3 8.08 9.57 10.50 10.21 11.11 12.14 12.12 13.28 12.58

Fe2O3tot 8.44 7.81 7.66 7.52 7.24 7.43 6.72 7.25 5.37

MnO 0.14 0.18 0.16 0.13 0.17 0.18 0.11 0.16 0.09

MgO 19.01 15.54 12.82 12.60 10.16 7.56 6.64 4.11 7.09

CaO 10.61 9.63 7.82 8.21 7.64 8.10 6.84 6.02 3.56

Na2O 1.92 3.01 1.99 1.33 2.29 2.03 2.54 3.60 3.16

K2O 1.57 1.75 3.23 5.65 4.18 4.89 5.44 5.68 5.32

P2O5 0.80 0.27 0.51 0.53 0.50 0.46 0.52 0.60 0.38

S 0.13 0.00 0.01 0.03 0.01 0.01 0.09 0.06 0.02

BaO 0.23 0.08 0.08 0.05 0.05 0.14 0.19 0.18 0.09

Cr2O3a 0.15 0.11 0.12 0.12 0.12 0.11 0.04 0.02 0.08

NiOa 0.09 0.05 0.04 0.03 0.03 0.03 \0.01 \0.01 0.02

LOI 4.20 6.42 4.37 2.06 3.67 3.84 3.07 0.56 3.38

O=S -0.06 0.00 -0.01 -0.01 -0.01 0.00 -0.04 -0.03 -0.01

Total 99.84 100.11 100.09 100.05 100.08 99.88 99.93 99.99 99.84

Nib nd 392 358 262 296 241 67 220 nd

Sc 20.0 18.0 19.4 22.4 20.0 22.9 18.7 15.8 11.8

V 87.9 47.8 117.7 99.9 126.5 130.0 105.0 120.8 59.7

Crc 772 568 872 733 901 827 221 123 587

Co 5.2 8.9 19.9 21.8 24.7 22.9 24.0 16.9 20.5

Cu 8.0 9.8 56.6 13.8 35.0 39.9 15.2 37.5 19.1

Rb 55.4 97.4 122 172 116 129 115 214 205

Sr 2,436 2,207 1,857 994 955 791 1,124 1,115 946

Y 24.6 21.6 15.3 14.1 16.0 15.2 13.2 24.6 14.0

Zr 180 203 91 86 112 95 106 205 124

Nb 7.5 8.8 4.7 3.4 5.1 4.4 6.5 15.3 8.8

Cs 7.5 2.4 11.1 7.0 71.6 12.1 5.5 5.9 6.0

Ba 3,027 1,713 1,499 1,316 1,341 1,955 1,862 1,464 1,382

La 82.4 59.1 26.4 23.8 25.4 23.1 29.7 44.0 33.1

Ce 169 125 55.9 48.8 53.1 46.5 59.3 87.9 64.1

Pr 19.5 15.2 6.8 5.8 6.3 5.4 6.7 10.0 6.9

Nd 77.7 60.9 27.7 24.2 25.2 21.9 26.6 39.9 25.8

Sm 13.5 10.6 5.2 4.8 5.0 4.4 5.0 7.8 4.6

Eu 3.3 2.6 1.4 1.3 1.3 1.2 1.1 2.0 1.2

Gd 9.7 7.6 4.1 4.1 4.2 3.8 3.9 6.4 3.5

Tb 1.1 0.9 0.5 0.5 0.6 0.5 0.5 0.8 0.5

Dy 5.6 5.0 3.0 3.0 3.2 3.0 2.8 4.7 2.7

Ho 0.9 0.9 0.5 0.6 0.6 0.6 0.5 0.9 0.5

Er 2.2 2.1 1.6 1.5 1.6 1.5 1.3 2.4 1.4

Tm 0.3 0.3 0.2 0.2 0.2 0.2 0.2 0.3 0.2

Yb 1.7 1.9 1.5 1.5 1.5 1.5 1.3 2.4 1.4

Lu 0.2 0.3 0.2 0.2 0.2 0.2 0.2 0.3 0.2

Hf 4.4 5.2 2.5 2.4 2.9 2.5 2.8 5.3 3.4

Ta 0.9 0.6 0.3 0.3 0.4 0.3 0.4 1.0 0.6

Th 22.5 14.2 7.2 9.5 7.4 7.9 10.5 24.8 12.3

U 7.4 5.1 2.2 3.9 2.6 3.0 3.1 9.1 2.0

(La/Yb)N 33.2 21.3 12.3 11.1 11.2 10.5 15.4 12.6 15.9


123

Table 1 continued

Sample 81–862 83–105 83–730 83–437* 83–404 83–634 83–258 83–803 83–745Locality 20 27 14 32 31 8 33 5 17

SiO2 60.70 62.69 63.15 63.54 65.29 65.38 65.44 66.37 66.80

TiO2 0.75 0.43 0.45 0.44 0.41 0.25 0.32 0.36 0.20

Al2O3 14.42 14.48 16.51 16.67 14.66 15.29 14.77 15.73 14.81

Fe2O3tot 6.11 3.54 3.89 4.12 3.34 2.44 2.52 3.08 1.77

MnO 0.15 0.12 0.11 0.14 0.04 0.07 0.03 0.08 0.10

MgO 2.21 2.15 0.98 1.42 1.75 2.12 1.18 1.38 0.55

CaO 2.39 4.67 3.51 3.61 0.93 1.93 2.78 2.46 2.59

Na2O 3.50 4.38 4.48 3.12 2.76 4.14 3.75 4.43 1.25

K2O 6.72 6.39 5.11 4.31 5.29 7.01 4.54 5.07 6.48

P2O5 0.55 0.31 0.22 0.5 0.30 0.21 0.20 0.20 0.14

S 0.05 0.01 0.00 nd 0.89 0.09 0.27 0.02 0.04

BaO 0.21 0.23 0.27 nd 0.11 0.17 0.15 0.17 0.26

Cr2O3a 0.02 0.01 \0.01 nd \0.01 0.01 \0.01 \0.01 \0.01

NiOa \0.01 \0.01 \0.01 nd \0.01 \0.01 \0.01 \0.01 \0.01

LOI 2.04 0.47 1.21 2.32 3.02 0.87 3.76 0.50 5.06

O=S -0.03 0.00 0.00 nd -0.44 -0.04 -0.14 -0.01 -0.02

Total 99.79 99.88 99.88 100.19 98.33 99.93 99.56 99.85 100.03

Nib 63 33 7 nd 20 41 11 32 3

Sc 14.4 6.7 7.5 6.1 6.0 6.5 5.5 5.5 5.1

V 113.5 72.7 57.7 66.5 59.7 53.3 46.0 48.5 27.9

Crc 141 53.2 16.4 24.1 29.6 125 20.2 44.0 9.7

Co 16.2 11.9 5.2 8.3 3.7 5.7 63.7 6.0 2.5

Cu 28.2 29.2 10.7 44.4 141.6 115.1 41.5 7.6 9.1

Rb 266 278 146 142.2 232 258 229 146 337

Sr 1,070 1,765 1,378 1,216 504 644 530 1,292 701

Y 28.9 15.8 22.0 16.0 13.5 12.5 10.9 14.1 11.0

Zr 218 294 207 207 185 295 167 173 317

Nb 15.3 13.8 11.0 13.2 12.9 12.4 8.7 9.7 12.1

Cs 8.1 5.9 5.9 12.4 8.5 3.7 13.7 2.6 20.5

Ba 1,767 2,087 2,091 387 1,125 1,591 1,071 1,292 2,435

La 51.7 87.2 40.1 88.5 66.9 26.1 48.2 40.5 29.0

Ce 105 185 79.4 165 126 51 92.6 76 53

Pr 11.7 19.7 8.8 16.8 13.0 5.3 9.7 8.0 5.3

Nd 46.4 71.4 33.9 59.4 45.4 19.6 34.8 28.8 18.4

Sm 8.9 10.5 6.3 8.5 6.9 3.5 5.3 4.8 3.1

Eu 2.1 2.7 1.7 1.9 1.6 1.1 1.3 1.3 0.8

Gd 7.5 6.5 5.3 5.7 4.4 2.9 3.7 3.7 2.5

Tb 1.0 0.7 0.7 0.6 0.5 0.4 0.4 0.5 0.3

Dy 5.4 3.5 4.2 3.3 2.7 2.4 2.2 2.6 2.0

Ho 1.0 0.6 0.8 0.6 0.5 0.5 0.4 0.5 0.4

Er 2.8 1.5 2.2 1.5 1.3 1.3 1.1 1.3 1.1

Tm 0.4 0.2 0.3 0.2 0.2 0.2 0.1 0.2 0.2

Yb 2.8 1.4 2.2 1.4 1.2 1.5 1.0 1.3 1.2

Lu 0.4 0.2 0.3 0.2 0.2 0.2 0.2 0.2 0.2

Hf 5.9 7.6 5.2 5.3 4.8 8.0 4.5 4.5 7.4

Ta 1.1 0.8 0.7 1.1 1.2 1.0 3.3 0.8 0.7

Th 19.9 124.2 11.7 24.7 35.1 28.1 20.0 16.5 15.7

U 5.0 25.0 3.6 13.7 12.8 11.2 9.8 4.1 5.5

(La/Yb)N 12.7 43.0 12.4 41.6 36.4 12.1 32.2 21.7 16.3


123

Table 1 continued

Sample 83–920 92–78 83–743 83–275 83–358 83–760 81–927 83–305 83–328Locality 28 7 18 33 30 19 24 29 29

SiO2 67.13 67.87 68.07 68.64 69.20 69.79 69.89 70.38 73.96

TiO2 0.32 0.37 0.21 0.34 0.28 0.20 0.24 0.30 0.25

Al2O3 15.57 15.92 15.01 14.22 13.69 15.27 15.34 13.26 13.2

Fe2O3tot 2.13 2.60 3.01 1.53 2.04 1.61 2.06 2.49 1.59

MnO 0.07 0.04 0.09 0.07 0.07 0.04 0.07 0.06 0.01

MgO 1.11 0.41 0.18 1.09 0.92 0.29 0.55 0.98 0.61

CaO 1.66 0.52 0.20 2.09 0.96 1.62 1.85 1.30 0.65

Na2O 3.68 3.93 1.57 2.71 2.75 4.62 4.83 3.44 1.66

K2O 5.45 6.26 10.23 6.68 5.07 5.54 3.88 5.10 5.42

P2O5 0.20 0.13 0.12 0.21 0.21 0.13 0.13 0.17 0.12

S 0.18 \0.01 0.00 0.36 0.78 0.00 0.05 0.38 0.44

BaO 0.21 0.32 0.30 nd 0.07 0.25 0.19 0.08 nd

Cr2O3a \0.01 \0.01 \0.01 nd \0.01 \0.01 \0.01 \0.01 nd

NiOa \0.01 \0.01 \0.01 nd \0.01 \0.01 \0.01 \0.01 nd

LOI 1.91 1.53 1.06 1.89 2.89 0.52 0.73 1.59 1.9

O=S -0.09 0.00 0.00 0.18 -0.39 0.00 -0.02 -0.19 0.22

Total 99.52 99.89 100.06 99.64 98.54 99.88 99.79 99.34 100.03

Nib 9 46 4 9 12 5 3 15 4

Sc 5.3 4.7 4.2 5.3 4.3 4.9 3.8 9.8 7.7

V 43.8 38.6 27.9 53.5 41.1 24.7 28.7 37.4 22.4

Crc 47.5 16.7 27.7 20.5 23.0 6.7 4.9 18.3 13.4

Co 3.7 3.4 1.5 4.0 5.7 3.9 3.5 4.1 4.8

Cu 15.8 6.0 6.4 603 579 6.4 8.8 2,120 590

Rb 197 228 416 267 305 207 121 158 528

Sr 1,106 1,479 478 816 578 791 1,364 468 55.9

Y 11.5 11.8 9.8 10.8 12.5 34.3 7.8 14.6 62.8

Zr 148 394 157 189 139 131 115 153 196

Nb 8.5 41.8 11.3 9.5 15.1 10.2 7.0 18.8 19.1

Cs 5.7 12.1 6.8 6.2 9.6 7.5 6.1 7.2 18.3

Ba 1,618 2,735 2,696 1,100 685 1,932 1,540 654 1,239

La 51.8 101.5 26.5 58.3 68.3 37.6 19.3 33.2 70.0

Ce 100.2 179.3 49 108.6 131.3 63 39 57 146.9

Pr 10.5 19.2 4.9 11.3 12.8 8.4 4.4 7.6 15.7

Nd 38.0 63.7 17.4 39.9 43.3 34.7 17.1 27.0 58.2

Sm 5.8 9.1 2.9 6.1 6.2 6.8 3.2 4.2 11.7

Eu 1.6 2.7 0.7 1.4 1.4 1.9 0.8 0.9 1.3

Gd 3.9 4.9 2.4 4.0 3.8 7.3 2.4 3.0 11.5

Tb 0.5 0.5 0.3 0.4 0.4 1.0 0.3 0.4 1.8

Dy 2.3 2.8 1.8 2.2 2.3 5.7 1.5 2.3 11.8

Ho 0.4 0.5 0.3 0.4 0.4 1.0 0.3 0.5 2.4

Er 1.2 1.3 0.9 0.9 1.1 2.6 0.7 1.4 6.8

Tm 0.2 0.2 0.1 0.1 0.2 0.3 0.1 0.2 1.0

Yb 1.0 1.5 1.0 0.9 1.2 2.1 0.7 1.4 6.4

Lu 0.2 0.2 0.2 0.1 0.2 0.3 0.1 0.2 0.9

Hf 4.1 10.9 4.6 4.8 4.0 4.0 3.6 4.6 6.2

Ta 0.7 2.6 0.7 0.8 1.2 0.6 0.5 1.7 1.7

Th 27.0 78.6 13.9 20.1 38.9 10.7 8.3 28.9 37.1

U 10.0 17.1 6.2 4.6 12.6 2.2 3.5 7.0 8.2


123

diameters between 10 and 200 m or as dykes up to 100 m

in width and 1,000 m in length. The intermediate to acidic

rocks are found mainly as stocks that vary in diameter from

\100 m to more than 1,200 m. The mean cross-sectional

area of the felsic stocks is 0.5 km2, which compares with

0.1 km2 for the smaller mafic and ultramafic intrusions.

Most of the intrusions are not zoned and show little vari-

ation in major elements, trace elements, or isotope com-

positions between margin and centre. However, some

intrusions do show chemical variation, for example, the

Machangqing intrusion (Number 8 in Fig. 1; Table 1)

varies from quartz monzonite to granite. Some of the most

evolved felsic members of the suite, including the Tong-

chang (No. 2 in Fig. 1) and Machangqing porphyries (No.

8 in Fig. 1) along the Red River shear system and the

Yulong porphyry copper ore field along the Batang-Litang

fault system (Nos. 26–33 in Fig. 1), are associated with

porphyry copper mineralization.

Petrography

Mineralogically the ultrabasic rocks contain more than

20 % phenocyrsts, comprising forsterite (4–10 %), diopside

(5–10 %) and phlogopite (2–6 %), in a cryptocrystalline

groundmass of diopside, phlogopite, sanidine, leucite and

olivine. Accessory minerals are magnetite and apatite. The

grain size of the phenocrysts varies between 0.2 and 3 mm.

The basic rocks can be divided into shonkinite, trac-

hybasalt, shoshonite and lamprophyre. The shonkinite

intrusions have \15 vol % phenocrysts, which include

olivine (1–2 %), augite (1–5 %), biotite (1–3 %) and

orthoclase (2–6 %), set in a groundmass consisting mainly of

alkali feldspar, diopside, olivine and biotite. The principal

accessory minerals are magnetite and apatite. The pheno-

crysts have grain sizes between 0.2 and 3 mm. The trachy-

basalts contain 10–20 % phenocrysts, which are composed

of pyroxene, olivine and biotite whereas the groundmass

consists of pyroxene, biotite, basic plagioclase and minor

alkali plagioclase. Accessory minerals include magnetite,

ilmenite and apatite. The phenocrysts have grain sizes of

0.3–3 mm. Shoshonite intrusions contain 10–15 % pheno-

crysts, which are composed of diopside, olivine and basic

plagioclase. The groundmass is composed of olivine, diop-

side and feldspar. Accessory minerals include magnetite and

apatite. The grain size of the phenocrysts is 0.2–4 mm. The

lamprophyres contain 25–30 % phenocrysts with grain size

ranging between 0.4 and 0.9 mm set in a finer groundmass.

The principal phenocrysts are phlogopite (18–30 %) and

diopside (2–10 %), and the groundmass is composed of

plagioclase (40–50 %), pyroxene (5–10 %) and biotite. The

principal accessory minerals are apatite and magnetite.

The intermediate rocks include syenite porphyry, latite

and trachyte. The syenite porphyries contains 20–30 %

phenocrysts of orthoclase (10–15 %), anorthoclase

(6–10 %), biotite (3–5 %) and amphibole (1–5 %), with a

grain size ranging from 0.1 to 4 mm, set in a cryptocrys-

talline groundmass of K-feldspar, albite and biotite. The

latites contain about 10 % phenocrysts of pyroxene, olivine

and sanidine in a cryptocrystalline groundmass composed

of pyroxene, plagioclase, sanidine and olivine. Accessory

minerals include magnetite, ilmenite and apatite. The

trachytes contain 5–10 %, 0.2–2-mm phenocryst of sani-

dine, biotite and minor plagioclase in a groundmass of

K-feldspar, plagioclase and biotite. The accessory minerals

are magnetite, sphene, zircon and apatite.

The felsic rocks consist of 20–40 % phenocrysts of

quartz, plagioclase, K-feldspar, biotite and amphibole set in

a fine-grained groundmass of K-feldspar, plagioclase,

quartz and biotite. The grain size of most phenocrysts is

between 0.1 and 5 mm, with rare grains up 12 mm in

length. The accessory minerals include magnetite, zircon,

sphene and apatite, with minor rutile in some samples.

Geochemistry and geochronology

The major and trace element chemistry for 27 new repre-

sentative samples of the Cenozoic high-potassium igneous

rocks of the Red River-Ailao Shan fault and its northern

extension are presented in Table 1. These data, together

with 25 analyses taken from the literature, are presented in

Table 1 continued

Sample 83–920 92–78 83–743 83–275 83–358 83–760 81–927 83–305 83–328Locality 28 7 18 33 30 19 24 29 29

(La/Yb)N 34.1 46.0 17.7 44.5 39.7 12.4 17.8 15.8 7.5

Major elements determined by XRF at Macquarie University by B.W. Chappell, except those marked ‘‘*’’ which were measured by wetchemistry at GIG–CAS, see ‘‘Appendix’’. Trace elements by LA–ICP–MS at ANU. nd = not determined. Subscript N denotes chondritenormalization (Sun and McDonough 1989)a Cr2O3 and NiO by XRFb Ni by solution ICP–MSc Cr by LA–ICP–MS. Localities given in Fig. 1


123

Figs. 2, 3, 4 and supplementary Table S1. The previously

published analyses were by wet chemistry and were carried

out at the Guangzhou Institute of Geochemistry of the

Chinese Academy of Science. Thirteen of the intrusions

were analysed at both the ANU and GIG, although the

samples analysed were not the same (i.e. different samples

from the same intrusion). Twenty-three samples from sev-

enteen intrusions were also analysed for radiogenic isotopes

at GIG. The intrusions are characterized by low TiO2

(\0.9 wt%), high alkali contents with K2O/Na2O [ 1 (for

most samples), and a wide range in both MgO

(0.1–19.0 wt%) and SiO2 (42.2–74.0 wt%). The forsterite

content of the olivines in samples from the mafic and

ultramafic members is high, and ranges between Fo93 and

Fo87 (Xu et al. 2001), consistent with derivation from a

mantle source. Most of the data plot in the shoshonite field

on a K2O versus SiO2 diagram (Fig. 2b) and, in some cases,

in the ultrapotassic field on plots of K2O versus Na2O

(Fig. 2c). The rocks are highly enriched in the REE and

LILE, with high chondrite normalized La/Yb [(La/Yb)N]

ratios (8–46). This combination of geochemical character-

istics is typical of shoshonitic rocks, which are defined by

Morrison (1980) as having high K2O/Na2O ([1.0 at 55 %

SiO2), high K2O ? Na2O ([5 %), high LILE enrichment,

low TiO2 and high but variable Al2O3 (14.0–19.0 %). The

Cenozoic high-potassium igneous rocks of eastern Tibetan

are therefore classified as shoshonites.

Major element vs silica plots are presented in Fig. 2a.

The total Fe2O3, CaO and MgO decrease with increasing

SiO2, but there is a distinct change in the gradient at SiO2

*55 % for all of these oxides except CaO. The data

clusters are not tight for either arm of the two trends. The

SiO2 versus MgO plot shows an interesting contrast of

behaviour with scattered data for the more mafic rocks, but

a relatively tight trend for the felsic ones. Reversals in

trends on Harker diagrams could be due to the onset of

crystallization of a new phase in a fractionating magma.

The decrease in total Fe2O3, CaO and MgO, accompanied

by an increase in total alkalies and Al2O3, could be

attributed to early fractionation of olivine, followed by

olivine ? pyroxene crystallization, and the reversal in the

total alkalies and Al2O3 trend attributed to the onset of

feldspar and phlogopite crystallization at *55 wt% SiO2.

However, the trace elements, with their more sensitive

distribution coefficients, are not consistent with fractional

crystallization. For instance, the trend of Cr against MgO

(Fig. 3b) is only consistent with fractional crystallization

for samples with MgO \ 7 % (SiO2 [ 55 %). Samples

with MgO [ 7 % show no evidence of a systematic

decrease in Cr with decreasing MgO as would be expected

if the geochemistry of these rocks were controlled by

olivine ? pyroxene ± chromite fractionation. The flat Cr

trend seen in Fig. 3b, for samples with MgO [ 7 %, is

unusual and cannot be explained by fractional crystalliza-

tion as will be discussed in greater detail later.

The patterns of the mantle-normalized trace and lesser

major elements for all samples, as shown in Fig. 4, are typical

of rocks from a convergent plate margin setting. They are

characterized by being highly enriched in LILE and in the

LREE, with pronounced negative Nb, Ta, and Ti anomalies,

negative P anomalies in most patterns, weak or absent Eu

anomalies, and positive Sr anomalies. What is not typical is

that the incompatible trace element patterns have similar

shapes and abundances over a large SiO2 range

(42.2–74.0 %). Notice in Fig. 4d that, although there is con-

siderable spread in the data, the field occupied by samples

with high MgO ([10 %) coincides almost exactly with the

field occupied by low MgO (\5 %) samples. The field for the

intermediate samples, those with MgO between 5 and 10 %,

overlaps only the lower incompatible element concentration

portion of the high and low MgO fields. An element pair that

does show a systematic variation with major element chem-

istry is Zr–Hf. These elements display a small negative

anomaly in the ultramafic samples (MgO [ 10 %) that is

weak or absent in the mafic and felsic samples (Fig. 4).

The isotope data for the eastern Tibet shoshonitic suite,

taken from the literature (Xu et al. 2001; Wang et al. 2001;

Zhu et al. 1992; Zhang and Xie 1997), define a broad field

but show no systematic variation with SiO2 or MgO; the

ultramafic, mafic and felsic rocks all show a similar range

in isotopic ratios (Figs. 5, 6). This is particularly true for Sr

and Nd isotopes. All of these rocks are characterized by

high 87Sr/86Sr (0.7050–0.7094) and low 143Nd/144Nd

(0.5123–0.5126) so that the data cluster in the enriched

quadrant in Fig. 5. 206Pb/204Pb varies between 18.5 and

19.2, 207Pb/204Pb between 15.6 and 15.7, and 208Pb/204Pb

between 38.6 and 40.0. Samples with high MgO

(MgO [ 10 %) tend to be less radiogenic that those with

low MgO (MgO \ 5 %), but there is considerable overlap

between the fields. The data lie between the MORB field

and the fields for upper and lower crustal Pb (Fig. 6).

Oxygen isotopes can be used to test for the presence of a

sedimentary component in the shoshonites source region

(Valley 2003). We have made preliminary measurements

of O isotopes in zircons from three eastern Tibetan sam-

ples. The results, which are listed in Table 2, vary between

6.0 and 7.4, with a mean of 6.6, which is well above the

mantle value of 5.3 and requires some sediment in the

source region (Valley 2003).

Published geochronology of the Red River shoshonitic

suite is limited and most are K–Ar dates, which are unre-

liable in many circumstances (Pan et al. 1990; Deng 1998;

Zhang and Xie 1997), although some intrusions have been

dated by the more reliable 40Ar/39Ar method (Wang et al.

2001; Chung et al. 1998). We rectified this situation by U–

Pb dating zircons from eighteen intrusions belonging to the


123

eastern Tibet suite by laser ablation inductively coupled

plasma mass spectrometry (Liang et al. 2007). The results,

and the methods used to obtain them, are summarized in

Fig. 1. They show that the shoshonitic suite along the Red

River fault zone crystallized at 35 ± 2.5 Ma, whereas

those along the Batang-Litang fault system formed some-

what earlier, between 37 and 41 Ma.

Discussion

Hypotheses for the origin of shoshonites

As noted in the introduction, the most widely held

hypothesis for the origin of shoshonites is that they form

by low degrees of melting of hydrous mantle that has

Fig. 2 a Variation in total Fe calculated as Fe2O3, CaO, MgO, total

alkali and Al2O3 against SiO2. b Plot of K2O against Na2O and b plot

of K2O against SiO2, showing the fields for ultrapotassic and

shoshonitic magmas. Data from Table 1 with additional data from

Zhu et al. (1992), Xie and Zhang (1995) and Zhang et al. (1998)


123

been metasomatically enriched in LILE and LREE (e.g.

Williams et al. 2004). The available experimental evi-

dence is consistent with this hypothesis. Wyllie and Se-

kine (1982), for example, have shown that melting of a

hydrous mantle can produce potassic magmas. Ebert and

Grove (2004) have demonstrated that Tibetan shoshonites

can coexist in equilibrium with the mantle. They found

that multiple saturation of olivine, clinopyroxene, apatite

and phlogopite occurs at 1.0 GPa, equivalent to a depth of

30 km and suggested that melting of a metasomatized

garnet peridotite, which may have been phlogopite bear-

ing, produced the Tibetan shoshonites. More evolved

magmas were attributed to fractionation of olivine and

clinopyroxene. However, the Tibetan Plateau has a

thickness ca.65 km, which requires the melting pressure

to have been at least 2.0 GPa, a pressure at which the

separating melt would only have been saturated in olivine

and clinopyroxene.

Other hypotheses that have been suggested to explain

the late Cenozoic high-potassic magmas of northern and

south-western Tibet include subduction of continental crust

(Deng 1998), convective removal of the mantle lithosphere

(Turner et al. 1996; Miller et al. 1999), extension along

strike-slip faults (Yin et al. 1995), slab break off (Miller

et al. 1999), and melting of metasomatized mantle and

sediments above a subducted slab (Arnaud et al. 1992).

Wang et al. (2001) suggested a metasomatized subconti-

nental lithospheric mantle source above continental crust

that was subducted during transpression along the Red

Fig. 4 Primitive mantle-normalized trace element patterns for rep-

resentative igneous rocks from the western Yunnan and eastern Tibet

for samples with MgO [ 10 % (a), samples with MgO between 10

and 5 % (b), and samples with MgO \ 5 % (c). The normalizing

values used are those of Sun and McDonough (1989). The patterns

from the different groupings are superimposed in Fig. 4d to show

their similarity. Blue MgO \ 5 %: orange MgO between 10 and 5 %:

red MgO [ 10 %

Fig. 3 Variations in Ni, Cr, Ce and Zr plotted against MgO


123

River-Ailao Shan-Batang-Lijiang fault system. There is no

consensus as to the origin of the Tibetan shoshonitic series.

Constraints

The mafic to ultramafic rocks, with MgO contents of up to

19 %, SiO2 contents as low as 42 %, and olivines with

forsterite contents around Fo90, have the essential charac-

teristics of a mantle-derived magma. These rocks could be

the product of melting of metasomatized mantle. The felsic

members of the suite, on the other hand, with SiO2 between

63 and 74 %, MgO between 0.18 and 2.12 %, and

phenocrysts of plagioclase, amphibole and quartz, cannot

possibly be in equilibrium with the mantle. If these felsic

magmas were derived by small degrees of partial melting

of a metasomatized mantle lherzolite, they would leave the

mantle in equilibrium with olivine and orthopyroxene.

Their SiO2 content would be buffered by these minerals

and should be *50 % or less, depending on the pressure

(Campbell and Nolan 1974); their Mg/Fe ratio would be

Fig. 5 Plot of 143Nd/144Nd against 87Sr/86Sr for igneous rocks from

west Yunnan and eastern Tibet. The fields for MORB, EMI, EMII and

GLOSS (global average subducted sediment (Plank and Langmuir

1998) are shown for comparative purposes. Data sources Xu et al.

(2001), Wang et al. (2001), Zhu et al. (1992) and Zhang and Xie

(1997). Comparative data for younger shoshonites from central and

western Tibet taken from Arnaud et al. (1992), Turner et al. (1996),

Miller et al. (1999) and Williams et al. (2004). Data fields from Ito

et al. (1987) and Hart (1988)

Fig. 6 Plots of 207Pb/204Pb against 208Pb/204Pb (a, c) and 206Pb/204Pb

against 208Pb/204Pb (b, d) for igneous rocks from western Yunnan and

eastern Tibet. The fields for MORB, EMI, EMII, HIMU, GLOSS,

Mariana ocean arc basalts and the northern hemisphere reference line

(NHRL) are shown for comparative purposes in a and b. The dashed

ellipse in c and d encloses the data from this study and is shown as a

horizontally shaded ellipse in a and b. Comparative data from sources

given in Fig. 5. Mariana data from a compilation provided by Jon

Woodhead. Data fields from Ito et al. (1987), Doe and Zartman (1979)

and Hart (1988). Other data source as for Fig. 5


123

controlled by the Mg/Fe of mantle olivine and orthopy-

roxene (Roeder and Emslie 1970), and they would not

contain quartz phenocrysts. Metasomatic enrichment of the

mantle source region in LILE and LREE may explain the

trace element and isotopic characteristics of the felsic

magmas but cannot account for their high SiO2, low Mg/Fe

or the presence of quartz. The quartz-saturated magmas

could be produced by extreme fractional crystallization of

mantle-derived magma but, if this were the case, the

amount of mafic magma would greatly exceed the amount

of felsic magma. The reverse is the case. The challenge is

to explain why the felsic magmas, which appear to have

formed by melting continental crust and the mafic–ultra-

mafic members, which are mantle melts, have the similar

trace element and radiogenic isotope signatures.

Another important consideration is the tectonic setting of

the Red River shoshonitic series. This shoshonitic suite,

unlike those from central and western Tibet, is not related to

subduction. It occurs in a restricted zone that is 1,500 km

long and 50 km wide, which closely follows the Red River-

Ailao Shan-Batang-Lijiang fault system and strikes at about

60� to the motion of the Indian plate. Furthermore, the field

evidence suggests that the shoshonitic suite is synchronous

with the fault system (Liang et al. 2007). There is no sug-

gestion of subduction in this region and therefore no justi-

fication for suggesting the Red River shoshonitic suite

formed by melting mantle during subduction.

Role of fractional crystallization

One possibility that must be considered is that the mafic and

ultramafic magmas formed by partial melting of a

metasomatized mantle and that the felsic magmas were

derived from these magmas by fractional crystallization.

Because the individual intrusions are small (average cross-

sectional area 0.5 km3), and widely spaced over an area that

is 1,500 km long and 50 km wide, the intrusion cannot

originate from a single evolving body at depth. However,

individual felsic magmas could have evolved from the mafic

magmas in the same region as represent by the sampled

mafic intrusions. Although considerable scatter is expected

in the data because the hypothesis requires fractionation to

occur in several genetically related but physically separated

systems, geochemical trends should still follow the path

predicted by fractional crystallization. Figure 3 shows two

compatible trace elements, Ni and Cr, for the eastern Tibetan

shoshonites. They show contrasting behaviours. Ni decrea-

ses with decreasing MgO as expected for a compatible ele-

ment controlled by fractional crystallization although the

rate of decrease is less than might be expected for strongly

compatible Ni. The concentration of Cr, on the other hand,

shows no systematic change with decreasing MgO for

samples with[7.0 % MgO. The absence of any decrease in

the Cr content of the eastern Tibetan shoshonitic suite, for

samples with MgO contents above 7 %, cannot be explained

by fractional crystallization but, as we will show, is consis-

tent with our continental crust melting model.

The incompatible element variations are also inconsis-

tent with fractional crystallization. In a fractionating system

they should increase systematically with decreasing MgO.

The eastern Tibetan data show no evidence of a systematic

increase in incompatible elements with decreasing MgO,

which is inconsistent with the observed trace element var-

iation being due to fractional crystallization.

Other observations that make fractional crystallization

an unlikely mechanism to explain the overall chemical

variation seen in the intrusions of the Red River-Ailao

Shan-Batang-Lijiang fault system are as follows:

1. The mafic–ultramafic intrusions are consistently smal-

ler than the felsic intrusions, which is the opposite of

what would be expected if they were related through

fractional crystallization.

2. The CaO trend lacks a kink that is characteristic of the

onset of plagioclase fractionation (Fig. 2a), an uncom-

fortable observation given the rock compositions.

3. The absence of a significant Eu anomaly and lack of

variation in the magnitude of the Sr anomaly with

increasing SiO2 for all rock types (Fig. 4) can be used to

rule out plagioclase as a significant fractionating phase.

Although there is limited chemical variation within

individual pipes that can be attributed to fractional crys-

tallization, especially for samples with MgO \ 7 %, the

overall variation between pipes from ultramafic to felsic

rocks cannot be explained by this process.

Table 2 d18O for zircons from Tibetan shoshonites. Numbers in

parentheses refer to localities at Fig. 1

Spot/sample/location d18O value s.e.m s.e.f.

634-2 7.50 0.13 0.33

634-6 7.28 0.08 0.31

634-11 6.04 0.17 0.34

634-7 5.97 0.08 0.31

83-634 Dali (8) 6.70 0.80

371-2 7.25 0.15 0.33

371-7 6.74 0.08 0.31

371-9 6.21 0.13 0.33

371-8 5.64 0.10 0.32

83-371 Duoxiasongduo 6.73 0.52

909-5 7.42 0.09 0.31

909-2 6.97 0.10 0.32

909-4 6.37 0.12 0.32

909-3 6.24 0.13 0.33

81-909 Gongjue (32) 6.53 0.39


123

A related explanation for the chemical diversity in the

eastern Tibetan high-potassium intrusions, especially Cr

versus MgO shown in Fig. 3b, is that it is due crystal

accumulation. Texturally, these rocks are not cumulates.

Moreover, if the rocks with 5–10 wt% MgO were cumu-

lates, one would expect distinctive patterns in Fig. 4,

namely higher compatible and lower incompatible ele-

ments. In fact, the incompatible trace element fields for

mafic and felsic rocks overlap. Accumulation of mantle

olivine phenocrysts or xenocrysts might contribute to the

formation of the ultramafic members of shoshonitic suite

by reducing their SiO2 and increasing their MgO contents.

Although the accumulation of mantle olivine might be

responsible for the transition of basic to ultrabasic, it

cannot explain the range of compositions from felsic to

ultrabasic. Crystal accumulation may have a second-order

influence on the chemistry of the eastern Tibetan high-

potassium intrusions but it is not the controlling process.

Magma mixing

The sub-parallel incompatible trace element patterns seen

in Fig. 4 could be due to magma mixing. This explanation

requires both end members to have similar incompatible

trace element patterns, which is unlikely, bearing in mind

that one end member has mantle affinities with

MgO = 19 % and SiO2 = 42 % and the other is crustal

with MgO = 0.2 and SiO2 = 74 %.

Role of assimilation

Another possible mechanism that might explain the shos-

honitic series is that they are mantle-derived melts that

have been variably contaminated by continental crust.

Distinguishing between assimilation of continental crust by

mantle melts and crustal melts that have interacted with the

mantle can be difficult if the incompatible trace element

contribution from the crust is small but is trivial if it is

large. Assimilation cannot raise the trace element content

of the contaminated melt above that, or even approach that,

of the contaminant, whereas if the contaminant is in the

source, this can readily be achieved at low or moderate

degrees of partial melting. The concentrations of highly

incompatible elements in the shoshonitic intrusions of

eastern Tibet are about twice the average for the conti-

nental crustal abundance (see Fig. 4). For example, Ce in

the shoshonite series varies between 39 and 185 ppm and

averages 91, compared with 48 in average continental crust

(Rudnick and Fountain 1995). Assimilation of average

continental crust would dilute Ce concentrations in the

high-Ce members of the shoshonitic series, which have Ce

concentrations as high as 185 (Table 1). Furthermore, even

if it is assumed that the contaminant had an anomalously

high incompatible element content (Th, U, Ce, etc.), the

amount of contamination required would still be high,

which would raise the SiO2 content of the mafic and

ultramafic members above the observed level. Moreover,

the increase in incompatible element concentrations

between mafic and felsic members, over a SiO2 range of

42–74 %, is less that 20 % and not systematic as would be

expected if the incompatible element concentrations were

controlled by assimilation (Fig. 3c, d).

Although the incompatible trace element characteristics

of the shoshonitic suite cannot be explained by assimilation

of continental crust, the geochemical characteristics of the

felsic melts are most consistent with a hypothesis that

involves melting of continental crust. Their high 87Sr/86Sr,

low 143Nd/144Nd, Pb isotopic characteristics, high SiO2,

low Mg/Fe, high LILE concentrations, LREE enrichment,

and pronounced negative Nb, Ta, and Ti, are all consistent

with a crustal source.

Continental crust melting and mantle interaction

hypothesis

Given the observations described above, we suggest that

the eastern Tibetan felsic shoshonitic intrusions were

principally the product of melting of continental crust.

They were intruded along the margins of giant Red River-

Ailao Shan-Batang-Lijiang fault system, which was trans-

pressional between 42 and 24 Ma (Ratschbacher et al.

1996). The close spatial relationship between the shos-

honitic suite and the Red River-Ailao Shan-Batang-Lijiang

fault system, and the correlation between the onset of

transpressional tectonics in this region and the start of

shoshonitic magmatism, are unlikely to be coincidences.

We suggest that the compressional component of the

transpressional movement along the Red River-Ailao and

Shan-Batang-Lijiang faults had a vertical component that

pushed one of the adjacent continental blocks into the

upper mantle (Wang et al. 2001). As noted earlier the

Moho for northern block is 5 to 10 km below than that for

the southern block. Geochemistry offers no special insight

as to the details of this process, but we suggest that the

most likely scenario is that southern block was partially

thrust below the northern block as shown in Fig. 7. This

would bring the upper surface of the lower block into direct

contact with higher temperature upper mantle. Heat con-

ducted from the mantle into the top of the down-thrust

crustal block would gradually heat it and raise its temper-

ature. In a different tectonic setting, a fragment of conti-

nental crust could be dragged into the upper mantle by the

subducting slab or delamination of the continental mantle

lithosphere. The essential requirement of our hypothesis is

that the crustal material is thrust deeply enough into the


123

upper mantle for the temperature of the adjacent mantle to

be sufficient to melt the down-thrust crust.

If a block of continental crust is pushed into the upper

mantle, the water content of that block will play a critical

role in determining how much, if any, melting occurs.

Sediments are the dominant rock type in the Himalayas,

and they can be expected to contain a moderate amount of

water held in hydrous phases such as clay minerals,

amphibole and zoisite. These minerals break down with

progressive burial; the water released either reacts with

adjacent minerals to form new hydrous minerals or is lost

from the system. The only hydrous minerals that are stable

at pressures above 2 GPa, the minimum pressure at the base

of the Himalayan crust, are the potassium-rich micas bio-

tite and phengite (Hermann and Green 2001; Hermann

2002). As a consequence, water can only be carried to the

base of the Himalayan crust in K-rich rocks, and the

amount of water that can be taken to these depths is pro-

portional to the amount of K in the transporting rock

(Patino-Douce and McCarthy 1998). At *2 GPa, biotite is

the dominant mica but it is gradually replaced by phengite

over a pressure range of *2–3 GPa. The pressure at which

the eastern Tibetan shoshonitic suite formed is uncertain

but, for our hypothesis to be valid, they must form below

the base of the Himalayan crust, that is at a pressure

[2 GPa. We have assumed 2.5 GPa where both biotite and

phengite are stable. At this pressure, the melting of conti-

nental crust will start at 700–800 �C and is controlled by

the breakdown of biotite and phengite according to the

reactions that have the general form (Hermann and Green

2001; Hermann 2002):

Phengite þ biotite þ clinopyroxene þ quartz

¼ garnet þ sanidine � kyanite � orthopyroxene

þ melt:

The order in which the minerals are consumed is

phengite then biotite followed by clinopyroxene. Garnet is

an important residual phase (20 ± 10 %) at all levels of

partial melting, and its fraction in the residue will increase

with partial melting. Quartz, clinopyroxene, sani-

dine ± kyanite and orthopyroxene will also be present but

their fraction in the residue will decrease as the melt

fraction increases (Hermann and Green 2001; Hermann

2002). At higher pressures, melting occurs at a higher

temperature, phengite replaces biotite and coesite replaces

quartz. At lower pressures, phengite is replaced by biotite

and muscovite, and melting starts at lower temperatures

with muscovite melting before biotite. If the composition

of the melted continental crust can be represented by

greywacke or pelite, the melts produced will be charac-

terized by high SiO2 (up to 75 %) and high alkalies, with

K2O [ Na2O (Hermann and Green 2001; Schmidt et al.

2004), which is typical of the felsic members of the shos-

honitic suite from eastern Tibet. The persistence of residual

garnet and sanidine, to high degrees of melting, has

important implications for the incompatible trace elements

geochemistry of these rocks. Abundant garnet in the resi-

due is required to explain their high LREE/HREE, and

residual sanidine will buffer the K2O content of the melt,

which may explain why the K2O of shoshonitic suite varies

between 4 and 7 % and does not change with SiO2 (Turner

et al. 1996). However, the amount of residual sanidine in

the mafic and ultramafic rocks cannot be large or they

would have negative Ba anomalies in their trace element

Fig. 7 Schematic tectonic model for the transpressional phase of

movement on the Red River-Ailao Shan-Batang-Lijiang fault system,

extensively modified from Wang et al. (2001). A vertical component

in the transpressional shear system pushes one of the crustal blocks

into the mantle where it starts to melt. a If the volume of melt

produced is large, it escapes to the surface with little interaction with

the overlying mantle to produce felsic shoshonites. b If the volume of

crustal melt is smaller, it percolates into the overlying mantle where it

reacts with olivine and orthopyroxene to produce mafic or ultramafic

members of the shoshonitic suite, depending on the extent of the

reaction, before eventually acquiring enough buoyancy to escape

from the mantle


123

patterns, which are not observed (Fig. 4). Negative Ba

anomalies are only seen in the felsic rocks.

The principal host for the LREE, Th and U in crustal

rocks is monazite, and the behaviour of these elements

during crustal melting is controlled by its solubility in the

melt unless all of the monazite is consumed during melting.

Zircon is relatively minor host for Th and U (Bea 1996),

and its presence or absence as a residual mineral will not

have a significant influence on the REE-Th-U content of

partial melts. Furthermore, as shown below, zircon is not a

residual mineral during partial melting to form the felsic

end members of the suite. Monazite has a preference for

the LREE and Th over U (Stepanov et al. 2012). Low-

temperature crustal melts formed by low degrees of partial

melting should leave residual monazite in their source, and

the complimentary partial melts should therefore be enri-

ched in U relative to the LREE and have Th/U ratios that

are lower than the protolith. High-temperature, high-degree

crustal melts, on the other hand, should not leave residual

monazite. The LREE, Th and U should behave as highly

incompatible elements, and their ratios in the melt will be

inherited from the source. The shoshonites from eastern

Tibet have high concentrations of LREE and have Th/U

ratios close to 4, the average value for the upper continental

crust (Taylor and McLennan 1985). This is consistent with

their formation by a high degree of partial melting of a

crustal metasediment and inconsistent with low-tempera-

ture, low-degree melting leaving residual monazite in their

source.

The Nb/Ta ratio of the eastern Tibet shoshonites pro-

vides further evidence that they formed by high degrees of

partial melting. Behaviour of Nb and Ta is controlled by

the minerals that host titanium (Stepanov and Hermann

2012). If the main Ti host in the restite is biotite/phengite,

Nb will behave as a moderately compatible element, and

the Nb/Ta ratio of the equilibrium partial melts will be

lower than that of the protolith (Stepanov and Hermann

2012). On the other hand, if Ti is hosted by rutile the Nb/Ta

ratio of partial melt will be higher than or similar to that of

the protolith, and much higher if it is hosted by titanite/

ilmenite. The Nb/Ta ratios of the Tibetan shoshonites from

Red River fault vary from 10 to 17, which compares with a

well-constrained value 12–13 for the continental crust

(Barth et al. 2000). The higher than crustal Nb/Ta ratios of

shoshonites indicate the presence of titanite/ilmenite or

rutile in their restite and the absence of mica. This is only

possible if the degree of melting is high enough to remove

all or most of the mica from the shoshonite source region.

The definitive features of shoshonites are their high K2O

content and high K2O/Na2O ratios, the later commonly

being between 1 and 2.5. Experimental studies by Hermann

and Spandler (2008) demonstrated that partial melting of

typical crustal rocks or metasediments, over a large range

of pressures and temperatures, produces peraluminous

granitic melts, which at crustal pressures, have K2O/Na2O

ratios close to unity. At higher pressure clinopyroxene,

with high jadeite content, becomes stabile and partial melts

become more potassic with K2O/Na2O ratios that increase

with increasing temperature and pressure. Hence, the high

K2O/Na2O ratios of shoshonites from eastern Tibet can be

explained by high-pressure melting of crustal rocks.

Additional evidence includes the absence of Eu anomalies

in the trace element patterns, which indicate absence of

plagioclase in restite, which in turn requires melting under

eclogite facies conditions. The presence of restitic eclogite

and granulite xenolith in mafic shoshonitic magmas from

Pamir, in the western high Himalayas (Hacker et al. 2005)

supports this hypothesis. Reaction between high-tempera-

ture magmas and the mantle can also create high-Na

pyroxene and lower K2O/Na2O ratios.

Our crustal melting-mantle reaction hypothesis is also

consistent with the variations of Ni and Cr with MgO

shown in Fig. 3a, b because our hypothesis requires the

highest MgO magmas to be the ones that have interacted

most thoroughly with the mantle. The principal reaction

produced by this interaction is the conversion of olivine to

orthopyroxene. Orthopyroxene contains high concentra-

tions of Cr especially at upper mantle pressures where it

charge balances Al3? in the tetrahedral site. Olivine, on the

other hand, contains negligible Cr. The reverse is true of

Ni. Olivine contains high concentrations of Ni and is its

major repository in the mantle, whereas orthopyroxene

contains only a few 100 ppm. As a consequence when the

quartz-saturated felsic melt, released by melting of conti-

nental crust, reacts progressively with the adjacent mantle,

Ni is released as olivine is replaced by orthopyroxene but

Cr, which would otherwise be released to the melt, is held

back by the crystallization of orthopyroxene. As a conse-

quence, the concentration of Ni increases with increasing

MgO, whereas Cr does not.

Three end-member cases need to be considered. First, if

a large block of continental crust is pushed deep into the

mantle so that it melts rapidly to produce a large volume of

felsic melt, it will have enough buoyancy to ascend rapidly

through the overlying mantle, probably in dykes, with

minimal interaction between the melt and the adjacent

mantle. If reaction does occur between the felsic melt and

the adjacent mantle, as discussed below, it will be confined

to a narrow zone in the adjacent mantle and it will have

little influence on the composition of the felsic melt

(Fig. 8a).

Second, if the volume of continental crust pushed into

the mantle is smaller than in the first case but still signif-

icant, or if the rate of melt production is lower, the initial

ascent of the magma is slower and the crustal melt will

percolate into the overlying mantle where it will react with


123

mantle as it rises. The crustal melt is quartz-saturated and

out of equilibrium with the olivine in the adjacent mantle.

It will react with this mineral to produce garnet and orth-

opyroxene, consuming melt and heat as it does so

(Campbell 1998; Yaxley and Green 1998). The precise

nature of these reactions is uncertain because there are no

published experiments on the interaction of a water

unsaturated crustal melt with peridotite at appropriate P

and T. However, experiments carried out by Rapp et al.

(2010), in which trondhjemite, tonalite and granodiorites

were reacted with mantle compositions in connection with

sanukitoid genesis, the layered eclogite-peridotite melting

experiments of Yaxley and Green (1998), and a study of

the system KAlSiO4–Mg2SiO4–SiO2–H2O by Sekine and

Wyllie (1982), give an indication of what to expect. Yaxley

and Green (1998) showed that if an eclogite layer, sand-

wiched between two peridotite layers, is heated above its

solidus, melting starts in the eclogite layer to produce a

dacitic melt with *65 % SiO2. This melt reacts with the

adjacent peridotite layer to produce orthopyroxene, garnet

and clinopyroxene. The composition of the melt produced

by melting continental crust will obviously depend on the

composition of the crust, the melt fraction and the condi-

tions of melting (P, T and water content). However, at

comparable melt fraction, the magmas produced by melt-

ing mica-bearing continental crust are expected to be richer

in SiO2 (up to 75 %), Na2O, K2O and H2O and poorer in

MgO, CaO and FeO than those produced by melting dry

eclogite and these differences will increase as the melt

fraction increases. As a consequence, interaction of the

crustal melt with peridotite is expected to produce more

orthopyroxene than observed in the eclogite sandwich

experiments (Yaxley and Green 1998). Garnet and clino-

pyroxene are also expected to form and phlogopite and/or

sanidine, depending on the amount of K2O and H2O in the

crustal melt and whether the percentage of melt in the

reaction zone remains high enough to prevent these min-

erals crystallizing.

As noted earlier, the amount of water in the system is

limited by the amount that can be carried into the mantle by

phengite and biotite and, as a consequence, the reaction

zone is expected to be water undersaturated. The Sekine

and Wyllie (1982) experiments were carried out under

water-saturated conditions, which limit the application of

their results to the problem under consideration. Water is a

stoichiometric constituent of phlogopite, which contains

two H2O molecules for each molecule of K2O, so the H2O

term is squared in the equilibrium equation, whereas the

K2O term is raised only to the power one. Water therefore

plays a more important role than potassium in determining

the stability field of phlogopite, which will obviously be

appreciably smaller in the water-undersaturated reaction

zone above the melting crust than in the water-saturated

experiments. Nevertheless, the available experimental data

show that phlogopite is stable over a wide range on tem-

peratures and some phlogopite can be expected to crys-

tallize in the reaction zone (unless the melt fraction

remains high), and this conclusion is consistent with the

occurrence of phlogopite phenocrysts in the mafic and

ultramafic intrusions. Crystallization of phlogopite, as the

rising crustal melt reacts progressively with the overlying

mantle, will be accompanied by a decrease in SiO2 as the

quartz-saturated melt reacts with olivine to form orthopy-

roxene, which explains why K2O decreases with decreasing

SiO2 at SiO2 \ 55 % (Fig. 2c).

Fig. 8 Sketch of two possible scenarios that could lead to the

formation of shoshonitic magmas if continental crust is pushed into

the mantle. a Production of felsic shoshonitic magma: heat conducted

from the mantle into the continental crust produces extensive melting

of the crust so that the accumulated melt has enough buoyancy to

escape through the overlying mantle in a dyke or pipe with little or no

interaction with the overlying mantle. b Production of mafic

shoshonitic magma: less melt is produced than in A or the melt

production rate is lower so that so the melt does not acquire enough

buoyancy to escape directly from the down-thrust continental crust

but percolates though the overlying mantle where it reacts, modifying

its composition. Eventually, its accumulating pool of melt in the

mantle acquires enough buoyancy to escape through a dyke or pipe


123

Provided the temperature in the reaction zone remains

high enough to prevent freezing, melt released from the

down-thrust crustal block will continue to react with the

adjacent peridotitic mantle until the accumulating melt

acquires enough buoyancy to escape from the mantle

(Fig. 8b). Here, it is important to remember that the tem-

perature of the mantle that reacts with the crustal melt will be

appreciably high than that of the crustal melt. The melt that

escapes from the mantle will be in equilibrium with olivine

and orthopyroxene. As a consequence, its SiO2 content will

be buffered by olivine and orthopyroxene, and its Ni content

and Mg/Fe ratio by equilibrium with olivine. Differences in

the MgO content will be controlled by relative contributions

of continental crust and mantle to the partial melting process

rather than by fractional crystallization. In this hypothesis,

the mantle controls the compatible elements in the mafic and

ultramafic melts and the incompatible elements, including

the radiogenic isotopes, are controlled by the crustal source.

It was stressed earlier that the reaction between the

crustal melt and the mantle consumes both melt and heat so

that the expected final volume of melt is appreciably less for

the mafic–ultramafic magmas, which were produced by

extensive reaction between the crustal melt and the mantle,

than for the essentially unreacted felsic magmas. As noted

in ‘‘Geological setting’’ section, the average cross-sectional

area of the mafic and ultramafic intrusion is 0.1 km2 com-

pared with 0.5 km2 for the felsic intrusions, which is con-

sistent with this interpretation. Another observation that is

consistent with the crustal melting hypothesis is the increase

in the size of the Zr–Hf anomalies with increasing MgO (see

Fig. 4). Our hypothesis requires less crustal melting in the

ultramafic members than in the felsic members. We suggest

that all of the zircon was consumed at the high degrees of

crustal melting required to form the felsic members,

whereas it was not at the lower degrees of melting required

to form the ultramafic members. As a consequence, Zr–Hf

anomalies are more pronounced in the ultramafic members

than they are in the felsic members. Because heat and melt

are consumed by the reaction between the continental crust

melt and the mantle, the mafic and ultramafic magmas

should be considered as the produce of this reaction and not

the result of melting of refertilized mantle.

Finally, if the down-thrust continental crust is low in

volume, or if the temperature of the mantle that reacts with

the ascending mantle is not high enough to prevent freez-

ing, all of the crustal melt will be consumed by the reaction

with the overlying mantle. The reaction will refertilize the

overlying mantle but no melt will escape to the surface.

Trace element modelling

Our hypothesis for the production of felsic magmas is

melting of a block of continental crust that has been pushed

into the mantle with minimal interaction with the sur-

rounding mantle. This is batch melting and can be modelled

using the batch melting equations. Modelling of the mafic

and ultramafic melts, which are thought to form by reaction

between the felsic melts and the adjacent mantle, is more

difficult because the reaction is not an equilibrium process.

However, equilibrium should be achieved on a local scale.

The principal effect on the trace element chemistry of the

reaction between the felsic melt and the adjacent mantle is

to dilute the incompatible trace elements and buffer the

compatible trace elements such as Ni and Cr with olivine

and pyroxene. The diluting effect of the mantle reaction has

been modelled by batch melting of mantle-continental crust

mixtures, combined in different proportions.

The results of the modelling are shown in Fig. 9.

Figure 9a shows the effect of varying the melt fraction F at

constant source composition and mineralogy, Fig. 9b the

effect of varying the amount of garnet in the residue, and

Fig. 9c the influence of changing the continent crust to

mantle ratio. The results show that changing F has a sig-

nificant influence on the highly incompatible elements

(LREE, Th and U) and Ce/Yb ratio but has little influence

on the HREE, changing the amount of garnet in the residue

influences the HREE but has little influence on the LREE,

and increasing the mantle to continental crust ratio in the

source region (as a proxy for the amount of mantle the

felsic melt reacts with) lowers the concentration of all

incompatible trace elements so that the trace element pat-

terns remain sub-parallel (i.e. element ratios remain

approximately constant). These calculations are not pre-

sented as unique solutions to the trace element geochem-

istry because the problem is not sufficiently constrained to

do so. For example, the observed HREE concentrations can

be modelled by (1) adjusting the amount of garnet in the

source residue, (2) varying the assumed HREE concentra-

tion of melted continental crust, or in the case of the mafic

and ultramafic melts, (3) varying the assumed mantle-

continental crust ratio in the residual source region. Simi-

larly the concentration of LREE and highly incompatible

elements can be varied by changing their assumed con-

centration in the melted crust or by changing F. Further-

more, there are no experiments available for continental

crust melt–peridotite interactions at an appropriate pres-

sure, similar to those carried out by Yaxley and Green

(1998) and Rapp et al. (1999) for eclogite melt–peridotite

interactions, which allow us to predict the residual miner-

alogy after the felsic melt has reacted with the adjacent

mantle. Nevertheless, the calculations do illustrate some

important principles. In particular, they show that (1) our

direct melting of continental crust model requires the

degree of partial melting to be high as might be expected if

low melting temperature continental crust is pushed into

relatively high-temperature mantle, (2) melting must occur


123

at sub-crustal depths in the absence of residual plagioclase,

(3) by making realistic assumptions, it is possible to model

the observed range of incompatible trace element concen-

trations and Ce/Yb ratios, (4) the ratio of adjacent elements

on the primitive mantle-normalized trace element diagrams

is inherited principally from the continental crust source

region and (5) the best results are obtained when the resi-

due has 20 ± 10 % garnet, in agreement with the experi-

mental work of Schmidt et al. (2004), who showed that a

garnet fraction persists in the residue of continental crust

anatexis to high degrees of melting. Given that the shosho-

nite compositions are highly variable, that these rocks spread

over hundreds of kilometres of fault strike length, and that

the composition of the crustal protolith is likely to be vari-

able, more sophisticated modelling is not warranted.

Major element trends

A feature of our model is that it explains why the major

element and Cr data follow two distinct trends, separated at

SiO2 *55 %. We suggest that the SiO2 \ 55 % samples are

in equilibrium with the mantle, whereas those with

SiO2 [ 55 % are not. Although the incompatible trace ele-

ment ratios and radiogenic isotopes in the low SiO2 samples

are inherited from the crustal block their SiO2 content, Mg/

Fe ratio and Ni and Cr contents are controlled by equilibrium

with mantle minerals. For melts with SiO2 \ 50 %, the

buffering minerals, which control these elements, are olivine

and orthopyroxene. There is no evidence that fractional

crystallization has played a role in their evolution. Samples

with SiO2 [ 55 % are not in equilibrium with the mantle.

Their geochemistry is controlled by the composition of the

source continental crustal block and by the conditions of

melting that block. Samples with SiO2 contents between 50

and 55 % lie between these end members. Melts in equilib-

rium with mantle olivine and orthopyroxene do not normally

have SiO2 contents much above 50 %. We suggest that

members of the shoshonitic suite, with SiO2 contents

between 50 and 55 % SiO2, formed where the anatectic melts

reacted with a limited volume of mantle and converted all of

the olivine in the reaction zone to orthopyroxene. The SiO2

content of the melts was therefore buffered by orthopyrox-

ene–quartz and will be higher than for the melts buffered by

olivine–orthopyroxene (Nicholls and Ringwood 1973). The

50–55 % SiO2 magmas lie on the extension of the

SiO2 \ 50 % trend for all oxides in Fig. 2, which is con-

sistent with this interpretation.

Heat source

Our hypothesis requires the heat needed to melt the con-

tinental crust to be extracted from the adjacent mantle. The

average cross-sectional area of the felsic intrusions is

0.5 km2. If it is assumed that they extend to a depth of

10 km and that the fraction of partial melting required for

their formation is 0.5, the volume of continental crust that

must be heated is 1016 cm3. If it is further assumed that the

continental crust source region must be heated to 850 �C to

achieve 50 % partial melting, the maximum amount of heat

required to raise the temperature of the source region is

4.5 9 1018 calories. Our assumed melting temperature is

consistent with an estimate of the emplacement

Fig. 9 Incompatible trace element models of batch melting of

average continental crust and continental crust-mantle mixtures.

a The influence of varying the melt fraction F at constant mineralogy

(10 % orthopyroxene, 10 % clinopryoxene and 20 % garnet) and at a

constant continental crust-mantle mass ratio. b The effect of varying

the amount of garnet in the residue at constant F and crust-mantle

ratio. c The effect of varying the continental crust to mantle ratio (as a

proxy for the extent of melt-mantle reaction, see text) at constant F

and residual mineralogy. Partition coefficients from Salters et al.

(2002) and average continental crust composition from McLennan

(2001)


123

temperature of 825 �C for the felsic intrusions by Liang

et al. (2006), based on Zr saturation. The volume of mantle

required to provide this heat is 1.2 9 1016 cm3. The

physical parameters used in these calculations for the

continental crust are as follows: specific heat =

0.32 cal gm-1 K-1, latent heat of fusion = 70 cal gm-1,

density = 3.0 gm cm3 and starting temperature = 600 �C;

and for the mantle, specific heat = 0.32 cal gm-1 K-1,

density = 3.3 gm cm3 and starting temperature =

1,200 �C. If the contact area between the crustal block and

mantle is 100 km2, the thickness of the mantle slab volume

required to provide the heat is 120 m and the thickness of

the slab of crust that must be heated is 100 m thick. That is

the total distance that heat must be conducted is 220 m.

The time scale for heat conduction can be calculated from

t ¼ X2=D

where t is time, X is the distance heat is conducted and D is

the conductivity. Taking D = 10-6 m2 s-1, the time

required is about 1,700 years. Reducing the contact area to

10 km2 increases the distance of heat conduction to 2,200 m

(1,200 ? 1,000) and the time required to 1.7 9 105 years.

These calculations show that if a block of continental crust is

pushed into hot mantle with a temperature of *1,200 �C (or

higher) at least the outer shell of the crustal block, within a

few thousand metres of the mantle, will melt on a geologi-

cally reasonable time scale to produce small volumes of

melt, similar to those observed in eastern Tibet.

Supporting evidence from xenoliths

Hacker et al. (2005) report both mafic and felsic crustal

xenoliths from Pamir, west of Tibet, which have been

brought to the surface by a 11 Ma potassic alkali magmas.

The felsic xenoliths are characterized by abundant garnet

plus omphacite, sanidine, quartz ± kyanite, with minor

rutile/titanite, scapolite, apatite and zircon, precisely the

mineral assemblage predicted from the crustal melting

hypothesis. Pressures and temperatures calculated from the

mineral assemblage vary between 2.0 and 3.0 GPa and

1,025–1,090 �C, respectively. Further, Hacker et al. (2005)

argue that the xenoliths are the residue of high degree (in one

case[40 %) dehydration partial melting of crustal material,

resulting from the breakdown of phengite and biotite. Weak

LREE enrichment and high concentration of HREE in the

xenoliths are consistent with this interpretation. Equilibrium

assemblages including mica species are conspicuously

absent although phengite inclusions in kyanite, and rare

biotite grains that are shielded by garnet, show that these

minerals were present in the rock prior to melting. Primary

melt inclusions in xenoliths of pelitic composition have

70 % SiO2, 16 % Al2O3, 2.1 % Na2O, 5.7 % K2O, 0.2 %

TiO2, with high LREE (Ce 140–240 ppm), Th, U and Ba and

low CaO, FeO and MgO (Chupin et al. 2006; Madyukov et al.

2011), a composition remarkably similar to the felsic

shoshonites from eastern Tibet.

Comparison with other hypotheses

The difference between our hypothesis and previous

hypotheses for the genesis of shoshonites is that melting in

our hypothesis is initiated by direct melting of continental

crust rather than by melting of mantle that was pre-con-

ditioned by metasomatized fluids derived from a subducted

slab or continental crust. The hypotheses that are most

similar to ours are those of Arnaud et al. (1992), Leslie

et al. (2009) and Prelevic et al. (2013) who argue for

involvement of subducted continental crust in the evolution

of shoshonites. In the Arnaud et al. (1992) model, the fluids

from the subducted slab fertilize the overlying mantle

wedge and crustal sediment, dragged down along the top of

the slab, are an important source of soluble incompatible

elements such as Th, U, Pb, Sr and the REE. They argue

that the mafic volcanic rocks formed by melting metaso-

matized mantle and that the associated felsic volcanic rocks

were produced by melting metasomatized sediments rather

than by direct melting of the down-thrust continental crust

and interaction of that melt with the adjacent mantle as in

our hypothesis. An important difference between our

hypothesis and those of Arnaud et al. (1992) and Leslie

et al. (2009) is that our hypothesis requires the shoshonites

to form by high degrees of partial melting (20–40 %) of a

crustal source, whereas other models require them to form

at low degrees of melting of a mantle source.

Prelevic et al. (2013) suggest that the widespread lam-

proites of the Mediterranean region, which are similar to

the more mafic end members of the shoshonite suite we

describe, are due to melting of a complex mixture of fly-

sch-type sediments and ultra-depleted oceanic-arc mantle

in a post-collisional environment to explain the mixed

ultra-depleted, ultra-enriched geochemical characters of the

lavas they studied. They attribute the mixed geochemical

characteristics of the Mediterranean lamproites to the

interaction of sediment derived melts with depleted mantle

and, in this respect, their hypothesis is similar to ours. The

Mediterranean lamproites have a number of notable trace

element characteristics that are also seen in the eastern

Tibetan shoshonites: a very high LREE content, highly

fractionated REE patterns, unfractionated crustal Th/U

ratios and LILE enrichment. The depleted characteristics of

the olivine and spinel in the melts described by Prelevic

et al. (2013), the enrichment in Li and P in the olivines and

their high Ni contents are all consistent with formation of

the lamproites through interaction of felsic melts with

mantle olivine. However, the Mediterranean lamproites are

found in a different tectonic setting to the eastern Tibetan


123

shoshonites, they do not include felsic members and it is

not suggested that they form at high degrees of partial

melting, an essential element of our hypothesis. We sug-

gest that both suites formed by high degrees of partial

melting of continental crustal material, followed by inter-

action with mantle peridotites.

Application to other shoshonites

The similarity in the incompatible trace elements and

radiogenic isotope ratios, over a range of SiO2 concentra-

tions from 42 to 74 %, is unique to the eastern Tibetan

shoshonitic suite and provides a rare insight into the origin

of this rock type. Furthermore, this shoshonitic suite is

formed in a well-constrained tectonic setting: adjacent to a

major transpressional fault separating two continental

blocks. Our hypothesis that the eastern Tibetan shoshonitic

suite formed by melting of down-thrust continental blocks

is not necessarily applicable to other shoshonites, which

display a smaller range in SiO2, especially those that form

in a different tectonic setting. However, the similarity

between the chemistry of the mafic and intermediate

members of the eastern Tibetan shoshonitic suite with

those from other locations suggests that this hypothesis

should also be considered for other occurrences.

Most shoshonites form in a convergent continental set-

ting, normally in association with subduction. In a number of

cases, shoshonitic magmatism can be linked to block faulting

and uplift. The significance of uplift is that it requires vertical

movement of one block relative to the other, the condition

required for our down-thrust hypothesis. Examples include

Papua New Guinea, Puerto Rico, the western USA (Norman

and Mertzman 1991), and the southern Andes between 16

and 26�S (Morrison 1980). In these cases, our hypothesis of

melting of down-thrust continental blocks may be directly

applicable. However, it cannot be applied to oceanic

shoshonites, such as those from the Mariana Islands, where

there are no continental blocks. Here, a different explanation

is required. Oceanic-arc shoshonites are normally attributed

to melting of mantle wedge that has been metasomatically

modified by fluids originating from subducted sediments

(Stern et al. 1988). An alternative is that the mantle wedge is

prepared for shoshonite generation by melts derived from the

sediments, but this requires unusually high temperatures at

the top the subducting slab.

The essential feature of the hypothesis is not the down

thrusting of continental blocks but the downward movement

of continental crust to the depth in the mantle where it starts

to melt. Geochemistry does not constrain the physics of the

process required to do this. The density of continental crust is

less than that of the mantle and it cannot descend to the

required depth unless it is pushed down by lighter material

above (e.g. a continental block) or it is dragged down by

being attached to a large mass of material that is denser than

the mantle, such as a subducting slab. The rarity of shosho-

nites shows that unusual conditions are required for their

formation. We suggest that the conditions required to push or

drag continental crust deep into the mantle occur rarely at

convergent plate margins.

Acknowledgments We thank the Tibet Geological Survey for

supporting our fieldwork. Liang Huaying thanks the Chinese Acad-

emy of Sciences and the China Scholarship Council for financing his

visit to the Australian National University. We are also indebted to

Professor Bruce Chappell who carried out the major element analyses

of the rocks. This work was co-supported by the ‘‘Strategic Priority

Research Program (B)’’ of the Chinese Academy of Sciences Grant

No. XDB03010302, and the Chinese NSF (41121002, 41272099,

41172080). Brendan Murphy, Sebastian Tappe, Cal Barnes, Steve

Eggins, Joerg Hermann and Hugh O’Neill are thanked for their

comments on the manuscript.

Appendix: Analytical methods of elements and isotopes

Major element XRF analyses of the whole rocks samples

were analysed by Prof. Bruce Chappell at Macquarie

University using the method of Norrish and Hutton (1969),

except for samples marked by star symbols (*), which were

analysed by wet chemical method at the Guangzhou

Institute of Geochemistry, Chinese Academy of Science

(GIG–CAS). The rare earth and trace elements concentra-

tions were measured by laser ICP–MS at the ANU, using

the method described by Campbell (2003). The glasses for

analyses were prepared by mixing finely powered rock with

lithium borate flux in the ratio 2:1, heating for 15 min at

1,200 �C, then quenched in water. The absolute concen-

tration of elements was determined by ratioing to Ca using

the NIST glass 610 as a standard. Analytical uncertainties

are ±1–2 % for major elements, and between ±2–10 %

(2r) for the trace elements, depending on the element.

Zircon U–Th–Pb dating was performed at the Research

School of Earth Sciences, the Australian National Uni-

versity following the procedure of Harris et al. (2004).

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the origin of shoshonites: new insights from the tertiary high-potassium intrusions of eastern tibet

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