Subduction metamorphism in the Himalayan1
ultrahigh-pressure Tso Morari massif: an integrated2
geodynamic and petrological modelling approach3
Richard M. Palin1,2*, Georg Reuber1, Richard W. White1, Boris J. P. Kaus1, and Owen M.4
Weller35
1Institute of Geosciences, Johannes-Gutenberg University of Mainz, 55128 Mainz, Germany6
2Department of Geology and Geological Engineering, Colorado School of Mines, Golden,7
80401, Colorado, USA8
3Department of Earth Sciences, University of Cambridge, Cambridge, CB2 3EQ, UK9
*Corresponding author: [email protected]
ABSTRACT11
The Tso Morari massif is one of only two regions where ultrahigh-pressure (UHP)12
metamorphism of subducted crust has been documented in the Himalayan Range. The13
tectonic evolution of the massif is enigmatic, as reported pressure estimates for peak14
metamorphism vary from ⇠2.4 GPa to ⇠4.8 GPa. This uncertainty is problematic for15
constructing large-scale numerical models of the early stages of India–Asia collision. To16
address this, we provide new constraints on the tectonothermal evolution of the massif via17
a combined geodynamic and petrological forward-modelling approach. A prograde-to-peak18
pressure–temperature–time (P–T–t) path has been derived from thermomechanical19
simulations tailored for Eocene subduction in the northwestern Himalaya. Phase20
equilibrium modelling performed along this P–T path has described the petrological21
evolution of felsic and mafic components of the massif crust, and shows that di↵erences in22
their fluid contents would have controlled the degree of metamorphic phase transformation23
in each during subduction. Our model predicts that peak P–T conditions of ⇠2.6–2.8 GPa24
and ⇠600–620 �C, representative of 90–100 km depth (assuming lithostatic pressure), could25
1
have been reached just ⇠3 Myr after the onset of subduction of continental crust. This26
P–T path and subduction duration correlate well with constraints reported for similar27
UHP eclogite in the Kaghan Valley, Pakistan Himalaya, suggesting that the northwest28
Himalaya contains dismembered remnants of what may have been a ⇠400-km long UHP29
terrane comparable in size to the Western Gneiss Region, Norway, and the Dabie-Sulu belt,30
China. A maximum overpressure of ⇠0.5 GPa was calculated in our simulations for a31
homogenous crust, although small-scale mechanical heterogeneities may produce32
overpressures that are larger in magnitude. Nonetheless, the extremely high pressures for33
peak metamorphism reported by some workers (up to 4.8 GPa) are unreliable owing to34
conventional thermobarometry having been performed on minerals that were likely not in35
equilibrium. Furthermore, diagnostic high-P mineral assemblages predicted to form in Tso36
Morari orthogneiss at peak metamorphism are absent from natural samples, which may37
reflect the widespread metastable preservation of lower-pressure assemblages in the felsic38
component of the crust during subduction. If common in such subducted continental39
terranes, this metastability calls into question the reliability of geodynamic simulations of40
orogenesis that are predicated on equilibrium metamorphism operating continuously41
throughout tectonic cycles.42
Keywords: Tso Morari massif; ultrahigh-pressure metamorphism; metastability; forward43
modelling; overpressure44
1 INTRODUCTION45
The Tso Morari massif, northwest India (Fig. 1), is one of only two examples of46
ultrahigh-pressure (UHP) metamorphism in the Himalayan Range and provides evidence47
for deep subduction of the Indian continental margin beneath Asia during the early Eocene48
(Epard and Steck, 2008). Constraining the petrological and metamorphic49
pressure–temperature–time (P–T–t) evolution of such eclogite-facies rocks is critical to50
understanding the geodynamic processes responsible for the burial and rapid exhumation51
of crustal material during collisional orogenesis (e.g. Warren et al., 2008). However, the52
2
tectonothermal evolution of the massif is controversial, with reported conditions of peak53
metamorphism varying from around the quartz–coesite transition (⇠2.4–2.8 GPa and54
⇠420–650 �C: St-Onge et al., 2013; Chatterjee and Jagoutz, 2015) to within the diamond55
stability field (⇠3.9–4.8 GPa and ⇠550–800 �C: Mukherjee et al., 2003; Wilke et al., 2015).56
This uncertainty hinders reliable large-scale tectonic reconstruction of the early stages of57
India–Asia collision.58
Thermomechanical modelling of heat transfer during plate-tectonic interactions allows59
calculation of theoretical P–T (–t) paths that a metamorphic rock may follow in a given60
geodynamic environment (Peacock, 1989). The P–T–t paths that these models produce61
may then be compared to the results of thermobarometric calculations performed on62
natural rocks. Forward modelling of slab-surface P–T conditions at subduction zones63
allows the influence of di↵erent geological variables on subduction zone thermal structures64
to be examined (Gerya et al., 2002), though the simulations are subject to significant65
uncertainties in the values of key parameters, including slab age, dip angle, and66
convergence velocity. The calculated evolutions of pressure and temperature with time67
provide tests of the geodynamic model within the uncertainty envelope.68
We have combined geodynamic and petrologic parameters specific to Eocene69
subduction of Indian-plate crust in the northwest Himalaya in simulations that provide70
new constraints on the prograde-to-peak P–T–t evolution of the Tso Morari massif.71
Petrological modelling was used to calculate bulk-rock properties for the main lithologies72
exposed in the massif at P–T conditions applicable to Phanerozoic subduction. These73
results provided inputs for a forward numerical model constrained by geodynamic criteria74
specific to India–Asia convergence, which predicted peak metamorphic conditions of75
⇠2.6–2.8 GPa and ⇠600–620 �C. Extremely high pressures up to 4.8 GPa inferred by some76
workers (Mukherjee et al., 2003; Wilke et al., 2015) appear to be geodynamically77
implausible. Further, diagnostic high-P metamorphic mineral assemblages predicted to78
form in felsic crust under equilibrium conditions are absent from the massif itself,79
suggesting widespread metastability during subduction. Our calculated P–T conditions are80
similar to those documented in the nearby UHP Kaghan Valley in Pakistan, which suggests81
3
that the northwestern Himalaya contains fragments of what may have been a coherent82
UHP terrane ⇠400 km long and ⇠150 km wide, comparable in size to the Western Gneiss83
Region, Norway, and the Dabie-Sulu belt, China.84
2 GEOLOGICAL BACKGROUND85
The Tso Morari massif, Ladakh Himalaya, is a structurally coherent block of thinned86
Indian continental margin (Mascle et al., 1994) exposed between Neotethyan sedimentary87
units of the Zanskar Range, and melange and ophiolite of the Indus Suture Zone (Fig. 1).88
The massif is ⇠100 km long, ⇠50 km wide, and ⇠7 km thick, and comprised primarily of89
quartzofeldspathic orthogneiss that hosts metre- to decametre-scale boudins of90
coesite-bearing mafic eclogite (De Sigoyer et al., 2004). Rare metasedimentary units also91
occur (Guillot et al., 1997). Field relations, major and trace element geochemistry, and92
U–Pb age data show that outcrops of Ordovician S-type granite exposed at Polokongka La,93
a low-strain portion of the massif (Fig. 1), represent relics of the orthogneiss’s protolith94
(Girard and Bussy, 1999). Eclogite boudins represent dismembered and metamorphosed95
mid-crustal doleritic dykes associated with Permian–Carboniferous Panjal Trap flood96
basalts exposed to the west (Fig. 1: Berthelsen, 1953; De Sigoyer et al., 2004).97
Geochronological and petrological work performed on both lithologies has identified98
multiple stages in the massif’s tectonothermal history. U–Pb zircon dating shows that the99
subduction of Indian continental crust initiated no later than c. 58–57 Ma (Guillot et al.,100
2003; Leech et al., 2005) and peak UHP conditions were reached at c. 51 Ma (St-Onge et101
al., 2013). These ages delimit a subduction duration of ⇠6–7 Myr, which correlates well102
with a ⇠7–9 Myr period inferred by Kaneko et al. (2003) for subduction of UHP eclogites103
in the nearby Kaghan Valley (Fig. 1). Initial and rapid exhumation to the base of the104
continental crust was followed by slower tectonic exhumation involving an105
amphibolite-facies metamorphic overprint at c. 45 Ma (U–Pb zircon, Leech et al., 2005;106
Th–Pb monazite, St-Onge et al., 2013). Ar–Ar data suggest exhumation through ⇠300–350107
�C at 30 Ma (De Sigoyer et al., 2000), although the timing of surface exposure is unknown.108
4
Despite the P–T conditions of retrograde amphibolite-facies overprinting in the crust109
being constrained at 1.0 ± 0.4 GPa and 650 ± 50 �C (e.g. De Sigoyer et al., 1997),110
estimates for peak pressure conditions fall into two distinct groups. Some workers suggest111
that subducted crust reached 80–100 km depth, inferred from P–T conditions of ⇠2.4–2.8112
GPa and ⇠420–650 �C around the quartz–coesite transition (St-Onge et al., 2013;113
Chatterjee and Jagoutz, 2015). However, others suggest subduction up to ⇠150 km depth,114
implied by putative P–T conditions of ⇠3.9–4.8 GPa and ⇠550–800 �C (Mukherjee et al.,115
2003; Wilke et al., 2015) well within the diamond stability field (Fig. 2). Notably, these116
discrepant P–T estimates were derived from rocks collected from the same roadside117
outcrop.118
Eclogite-facies metamorphism is recorded in mafic boudins by the assemblage garnet,119
omphacite, phengite, rutile, sodic-calcic/sodic amphibole, and quartz or coesite (De Sigoyer120
et al., 1997; Epard and Steck, 2008). Diamond has not been identified in the region. The121
host orthogneiss records negligible petrological evidence of UHP metamorphism, being122
characterised by the amphibolite-facies assemblage quartz, K-feldspar, plagioclase,123
muscovite, biotite, ilmenite, and garnet, with incipient partial melting (De Sigoyer et al.,124
2004; St-Onge et al., 2013). Nonetheless, Girard and Bussy (1999) and De Sigoyer et al.125
(2004) reported evidence for rare prograde reaction between magmatic plagioclase and126
biotite in the granite protolith to form Ca-rich garnet, kyanite, phengite, zoisite, and rutile.127
This assemblage occurs in some other high-P orthogneisses worldwide (e.g. in the Gran128
Paradiso and Monte Rosa Alpine terranes: Dal Piaz and Lombardo, 1986), and129
demonstrates that the Polokongka La granite protolith experienced metamorphic P–T130
conditions identical to the metabasites. Retrograde, mid-crustal overprinting of eclogite is131
recorded by a transition from garnet- and omphacite-dominated assemblages in boudin132
cores to an amphibolite-facies rim assemblage of quartz, plagioclase, calcic amphibole,133
epidote/clinozoisite, sphene, and biotite (De Sigoyer et al., 2000; Palin et al., 2014).134
5
3 INTEGRATED FORWARD MODELLING135
New constraints on the evolution of the Tso Morari massif have been obtained here via a136
three-stage forward-modelling approach that closely integrates geodynamic and137
petrological information. Firstly, phase diagram modelling was used to calculate the fluid138
contents and bulk-rock densities of stable assemblages that might have formed in139
metamorphosed felsic and mafic massif lithologies during subduction. Bulk-rock140
compositions for protoliths of the massif’s orthogneiss (Polokongka La granite) and eclogite141
(Panjal Trap basalt) were used instead of those for the metamorphic products themselves,142
owing to (1) uncertainty surrounding the relative timing of mineralogical transformations143
within the subduction–exhumation cycle, and (2) to avoid the e↵ects of open-system144
amphibolite-facies retrograde metamorphic overprinting on peak UHP assemblages. This145
approach provides an alternative framework with which to predict the petrological changes146
that occur during subduction of crustal materials. These bulk-rock petrophysical estimates147
were incorporated into a two-dimensional thermomechanical model specific to Eocene148
subduction in the northwest Himalaya, which generated P–T–t paths tailored for the149
region. The phase diagrams calculated previously were then examined within the context of150
these calculated P–T paths in order to determine the changes in mineral assemblage that151
might have occurred during subduction and metamorphism of Tso Morari massif crust.152
Subduction-zone thermal structures are primarily controlled by plate convergence rate,153
and the age and dip angle of the descending slab (Kirby et al., 1991). Changes in the154
metamorphic mineral assemblages of constituent rocks determine subducted crust density.155
Most studies of UHP terranes worldwide focus on mafic components due to their retaining156
detailed P–T (–t) information about the subduction–exhumation cycle; however, their low157
volume proportion (⇠1–2%) in upper continental crust means that they make no158
significant contribution to bulk-terrane petrophysical properties (Peterman et al., 2009).159
While experiments show that felsic crust can transform to (U)HP minerals such as jadeite,160
kyanite, and/or coesite, natural examples are rare, implying that most such subducted161
material (1) does not transform, (2) transforms and reverts to low-P assemblages during162
exhumation or later metamorphic overprinting, or (3) transforms at depth but is not163
6
exhumed. Many studies of collisional terranes worldwide support the first option (Carswell164
et al., 1986), although high-P felsic granulite that represents overprinted eclogite has been165
documented from a number of regional metamorphic terranes worldwide (O’Brien et al.,166
2003). The limited ability of some crustal rocks to re-equilibrate during subduction and/or167
exhumation is attributed to the lack of a critical catalyst such as fluid or deformation (e.g.168
Rubie, 1986). As such, phase transformations cannot be assumed to occur a priori, even169
though geodynamic models of collisional orogenesis often consider this to be so (Gerya and170
Stockhert, 2006; Warren et al., 2008). Our thermomechanical simulations thus consider two171
end-member scenarios: non-transformation (metastability) and complete transformation172
(equilibrium) of Tso Morari massif lithologies.173
3.1 Metamorphic phase equilibrium modelling174
All phase diagrams were constructed using thermocalc v3.40i (Powell et al., 1988) and175
the internally consistent thermodynamic dataset ds55 (Holland and Powell, 1998; updated176
August 2004) in the Na2
O–CaO–K2
O–FeO–MgO–Al2
O3
–SiO2
–H2
O–TiO2
–O2
177
(NCKFMASHTO) compositional system. The activity–composition relations used for178
solid-solution phases are listed in the Supplementary Information. Bulk-rock compositions179
used for modelling were converted from weight-percent oxide values reported in each180
original study (Table S1) to molar proportions of oxides (Table 1), with the water content181
of each protolith derived from reported XRF loss-on-ignition contents. Fractionation of the182
Panjal Trap basalt/eclogite bulk composition owing to sequestration of cations into garnet183
porphyroblasts during prograde metamorphism was also implemented using the rbi code in184
thermocalc (see Supplementary Information). Parameter maps showing calculated185
bulk-rock densities and free H2
O contents for both protoliths (assuming complete186
metamorphic transformation) at 300–800 �C and 0.2–5.0 GPa are shown in Fig. 3. Full187
phase diagrams showing the equilibrium assemblages for each rock type are shown in Figs188
S1–4. Uncertainties on the position of calculated assemblage-field boundaries in P–T space189
are up to ±1 kbar and ±50 �C at 2� (Powell and Holland 2008; Palin et al., 2016).190
In order to e↵ectively integrate our petrological modelling results into the geodynamic191
7
simulations described below, we assumed that metamorphism of the massif as a whole192
occurred in a closed-system environment. Although fluid influx from rocks external to the193
massif crust may have occurred during subduction, the absence of constraints on the194
timing and/or amount of fluid infiltration precluded its consideration. Nonetheless, we195
qualitatively examine the e↵ects of open-system processes by examining how fluid released196
from each protolith during devolatilisation could interact with the other. This massif-wide197
closed-system scenario represents the simplest interpretation of its evolution, and so defines198
one end-member of a possible hydrodynamic spectrum that can be explored in greater199
detail if and when constraints on fluid migration patterns become available.200
3.2 Geodynamic numerical modelling201
The numerical code MVEP2 employed for geodynamic numerical modelling is described in202
detail in Kaus (2010) and Thielmann and Kaus (2012), and can be downloaded from203
https://bitbucket.org/bkaus/mvep2. A brief summary is given here, alongside a description204
of the model parameters. However, as P–T–t paths calculated via such modelling are205
complex functions of the input parameters, sensitivity testing allowed an assessment of the206
reliability of our main results according to geologically realistic variation in these values.207
3.2.1 Governing equations and numerical approach208
The incompressible Stokes equations are defined as follows:209
@vi
@xi= 0 (1)
�@P
@xi+
@⌧ij
@xj= ⇢gi (2)
⇢cp
✓@T
@t
+ vi@T
@xi
◆=
@
@xi
✓k
@T
@xi
◆+ ⌧ij ✏ij � ↵T⇢gvz (3)
where vi is velocity, P is pressure, ⌧ ij is deviatoric stress, ⇢ is density, gi is gravitational210
8
acceleration, cp is heat capacity, T is temperature, k is thermal conductivity, ✏ij is strain211
rate, and ↵ is thermal expansivity. Both shear heating and adiabatic (de)compression were212
incorporated into the energy equation. Density of the upper crust in models considering213
mineral transformation varied as a function of P and T according to the values calculated214
via petrological modelling (Fig. 3).215
A viscoplastic constitutive relationship was used:216
⌧ij = 2⌘e↵ ✏ij (4)
⌘e↵ = min (⌘disl , ⌘pl) (5)
where the e↵ective (e↵ ) viscosity is the minimum of the dislocation (disl) and plastic (pl)217
viscosities, which are defined as:218
⌘pl =�yield
2✏II(6)
⌘disl = A
� 1n✏II
(1�n)nexp
✓E + PV
nRT
◆(7)
Here, �yield [= P ·sin(�) + c·cos(�)] is the yield stress as a function of the friction219
coe�cient, pressure, and cohesion, which followed a Drucker–Prager relationship (Drucker220
and Prager, 1952), ✏II [= (0.5·✏ij2)0.5] represents the second invariant of the strain rate221
tensor, A is a pre-exponential material parameter, n the power law coe�cient, E the222
activation energy, and V the activation volume. These equations were solved with the223
MATLAB-based, two-dimensional, thermo-mechanical finite-element code MVEP2 using224
Q1
P0
elements. Additionally, we applied a marker-and-cell approach to advect markers225
with the material properties above a Lagrangian mesh. Remeshing was applied every226
timestep to prevent the elements from becoming too distorted.227
9
3.2.2 Model setup and extraction of results228
Simulations considered a 4000-km wide and 660-km deep domain that used 801 ⇥ 201229
elements along the horizontal and vertical axes, respectively (Fig. 4). This produced a230
resolution in the subduction zone area of approximately 1 km ⇥ 1.5 km. Boundary231
conditions comprised a free surface (no stress) at the top, and free slip at the bottom and232
sides. Compression of the crust used a pushing box at the right-hand side of the domain233
with rates defined by various plate velocity scenarios (see below). Thermal boundaries were234
isothermal at the top and the bottom, and flux-free at the sides. In addition to shear and235
adiabatic heating, a geotherm of 9 �C/km was used for the crust and lithospheric mantle,236
and 0.5 �C/km was used for the aesthenospheric mantle. This produced a maximum237
temperature of 1680 �C at 660 km and ensured that the lithosphere was always colder than238
1350 �C.239
A weak zone was incorporated into the lithosphere in order to initiate subduction,240
which had a lower friction coe�cient and viscosity than its surroundings. We applied a241
constant sedimentation rate of 1 mm/yr, and a constant erosion rate of 5 mm/yr if the242
topography exceeded 2 km in elevation. After each simulation, we implemented a 4th-order243
Runge–Kutta scheme (Press, 1992) to advect markers backwards in time and extract P–T244
information. This provided the opportunity to examine the total range and variability of245
any model parameter over any spatial area.246
3.2.3 Density247
Density changes were investigated using two end-member scenarios: full transformation248
(i.e. equilibrium metamorphism) and non-transformation (i.e. metastability) of the Tso249
Morari massif upper crust. For the former, bulk-rock densities for assemblages in each250
metamorphosed protolith were calculated between 0.2–5.0 GPa and 300–800 �C (Fig. 3);251
however, as field studies estimate that mafic eclogite represents no more than ⇠1–2% total252
volume of the Tso Morari massif, the terrane density at any P–T condition was considered253
as that of the metamorphosed Polokongka La granite protolith only. For254
10
non-transformation, the massif was assigned a fixed density of 2700 kg/m3.255
3.2.4 Geometry of the subducted slab256
Numerous possible slab geometries for subduction of Indian crust in the Ladakh Himalaya257
were outlined by Leech et al. (2005). We used geometric parameters from their preferred258
configuration (“Model 1”), which had a geologically realistic dip angle that increased with259
depth and a curved slab surface with a bending radius of 350 km. Under these conditions,260
the slab dips at ⇠6� at the trench and ⇠40� at ⇠100 km depth (the approximate HP–UHP261
transition: Fig. 2). Their model assumed an initial mid-crustal intrusive depth of 15 km262
for the massif’s mafic dyke and granite protoliths, which also we adopted here.263
3.2.5 Subduction velocity and duration264
Two subduction velocity profiles were considered in our geodynamic modelling: (1) a265
constant 7 cm/yr velocity, as proposed by Leech et al. (2005) in their preferred Model 1266
setup, and (2) a time-dependent velocity profile simulating the reported deceleration of the267
Indian indenter immediately following initial collision with Asia. This considered a 10268
cm/yr velocity for the first 2 Myr following the onset of crustal subduction and 4 cm/yr269
afterwards (modified from Guillot et al., 2003). These two profiles were initially combined270
with both mineralogical transformation scenarios described above to produce four271
end-member reference models (M1–4), as detailed in Table 2. All combinations used the272
above-described geometric setup and produced very similar prograde-to-peak P–T–t paths273
(Fig. S5); however, as a time-dependent deceleration velocity profile is more geologically274
realistic, we focus our discussion below on the results of these simulations (M3–M4).275
3.2.6 Sensitivity testing276
The sensitivity of our P–T–t results to the model setup conditions was investigated by277
re-running simulations M1–M4 whilst individually varying each of 12 key model278
parameters. These variables comprised: bulk-rock composition, convergence rate, upper279
11
and lower crustal flow laws, weak zone angle, sedimentation rate, erosion rate, friction280
angle of the upper and lower crust, heat production, heat capacity, and281
temperature-dependent conductivity (Table S4). The influence of bulk composition was282
investigated by recalculating densities for a fully-hydrated Polokongka La granite protolith.283
These were similar to those produced at low pressure (<15 kbar) using the actual284
XRF-derived water content, but lower at higher pressures owing to the stabilisation of285
white mica in place of K-feldspar and jadeite.286
3.3 Results287
3.3.1 Main simulations288
Figure 2 shows P–T–t paths obtained from simulations M3 and M4 (deceleration with or289
without phase transformations). Upper crustal elements tracking the leading edge of the290
Indian plate margin initiated at ⇠0.4–0.5 GPa and ⇠160–190 �C at t = 0 Myr,291
representative of an initial depth in the crust of ⇠10–15 km.292
The complete transformation of massif components to high-P assemblages during293
subduction (M4) predicts ⇠1.2 GPa and ⇠240 �C to be reached after ⇠1 Myr of294
subduction, with a significant P–T increase up to ⇠2.4 GPa and ⇠550 �C in the next 2295
Myr (Fig. 2). Continued subduction until t = 7 Myr—the upper limit of subduction296
duration—produced only a small increase in P–T conditions to ⇠2.9 GPa and ⇠620 �C,297
with the tracked elements representing the leading edge of the Indian plate terminating298
just above the quartz–coesite transition (Fig. 2).299
The modelled P–T path for non-transformation (M3) was very similar to that for300
transformation, though subduction was slightly slower (Fig. 2). Near-UHP conditions of301
⇠2.5 GPa and ⇠425–575 �C were reached after ⇠4 Myr of simulation time, with similar302
peak P–T conditions of ⇠2.9 GPa and ⇠600 �C reached at t = 5.5 Myr (Fig. 2).303
Continuation of this model until t = 7 Myr did not significantly increase this peak P–T304
condition, as the low density of the crust causes it to stall and thicken at ⇠100 km.305
Notably, both modelled paths lie within the range of modern-day subduction-zone slab-top306
12
P–T profiles reported by Syracuse et al. (2010).307
The equilibrium mineral assemblages in both protoliths along the full transformation308
(M4) P–T path are shown in Fig. 5. Low-grade (<1.7 GPa, <350 �C) Panjal Trap309
metabasalt assemblages are dominated by glaucophane, chlorite, and epidote, with lesser310
actinolite, omphacite, muscovite, lawsonite, and sphene (Fig. 5a). Both garnet and rutile311
stabilise at higher grade at the expense of lawsonite, actinolite, and glaucophane. No free312
H2
O occurs at low/medium-grade conditions, as any produced via the breakdown of313
chlorite and epidote would be incorporated into other newly-formed hydrous phases.314
Excess fluid is not generated until relatively late in the subduction history at ⇠2.3 GPa315
and ⇠520 �C (Figs 3 and 5a). The UHP eclogite-facies assemblage stable at peak P–T316
conditions (⇠2.8 GPa and ⇠620 �C) is dominated by garnet (⇠30%) and omphacite317
(⇠40%), with talc, coesite, muscovite, lawsonite, rutile, and H2
O comprising 1–7% each318
(Fig. 5a), which closely matches assemblages in relatively fresh mafic eclogite from the319
massif (Epard and Steck, 2008). Putative continued subduction to ⇠4.8 GPa (dashed320
arrow on Figs 2 and 3), as suggested by Mukherjee et al. (2003) and Wilke et al. (2015),321
would involve the loss of talc and lawsonite, a minor increase in the proportions of garnet,322
omphacite, and coesite, and continued devolatilisation (Fig. 5a).323
The Polokongka La granite exhibits a notably simpler evolution during subduction.324
Under equilibrium conditions, the protolith assemblage K-feldspar, quartz, plagioclase,325
biotite, and muscovite transforms to jadeite/omphacite (⇠18%), muscovite (⇠13%),326
K-feldspar (⇠23%), quartz (⇠45%), garnet (⇠1%), rutile (<1%), and kyanite (<1%) at327
relatively low-grade conditions (<350 �C), with virtually no further changes predicted up328
to UHP conditions except for the replacement of quartz by coesite (Fig. 5b). The low H2
O329
content of the protolith (0.57 wt%: Table 1) keeps the rock fluid-undersaturated for the330
entire prograde P–T path, with water-saturated conditions restricted to low-P331
amphibolite-facies conditions (Fig. 3). Thus, prograde metamorphism would have332
proceeded under fluid-absent conditions where reaction kinetics are sluggish (Rubie, 1986).333
13
3.3.2 Additional simulations334
Sensitivity testing of models M1–M4 was achieved by running thirty-eight additional335
simulations that allowed the main e↵ect of each parameter on the style of subduction to be336
individually assessed. These results are shown in Figure 6 alongside our four reference337
models (bold red arrows). The evolution of each sensitivity test as a function of simulation338
time is shown in Fig. S6 and each test’s parameter combination is listed in Table S5.339
A lower density calculated for a fully hydrated Polokongka La granite protolith340
produced an increased buoyancy force that resulted in lower overall peak pressure341
conditions than M1–M4, as shown in Fig. 6 by the calculated mean P–T conditions of the342
entire subducted upper crust (“Hydrated crust” path). A greater plate velocity for initial343
collision (15 cm/yr; cf. Jagoutz et al., 2015) resulted in faster and steeper subduction, and344
thus calculated conditions at any equivalent simulation time relative to M1–M4 were at345
higher pressures (⇠0.5–0.8 GPa greater), but similar temperatures (Fig. 6: “High plate346
velocity” path). Nevertheless, since the upper crust is typically not subducted deeper than347
the lithosphere–asthenosphere boundary (Marquardt and Miyagi, 2015), the mean P–T348
conditions for upper crustal units was similar to M1–M4, albeit reached 1–2 Myr earlier349
due to more rapid convergence (Fig. S6). An increased weak zone angle up to 40� at the350
trench produced exceptionally steep and cold subduction through the HP–UHP transition,351
with subducted crust consistently entering into the Forbidden Zone (Fig. 6: “High weak352
zone angle” path).353
Varying sedimentation rate had no e↵ect on the subduction style, though high erosion354
rates led to a slightly shallower subduction angle. A higher friction angle for the upper355
crust produced smaller peak pressures, while a higher friction angle for the lower crust356
produced similar results to M1–M4 (Table S4). Heat production did not influence the357
subduction style in our simulations, although a higher heat capacity caused slightly358
shallower subduction and thus lower peak pressures than our main models (Fig. 6).359
Finally, implementing temperature-dependent conductivity with a non-linear crustal360
geotherm (Table S5 and Fig. S8) had the main e↵ect of increasing the mean temperature361
of subducted upper crust by approximately ⇠80–100 �C (Table S4), although calculated362
14
pressures at any given simulation time were very similar to M1–M4, and the upper crust363
mostly failed to reach pressures above ⇠3 GPa (Fig. 6: “Temperature-dependent364
conductivity” path).365
4 DISCUSSION366
4.1 Reliability of calculated P–T–t paths367
We designed an integrated geodynamic and petrological model to constrain the368
prograde-to-peak P–T–t evolution of Tso Morari massif crust during the initial stages of369
India–Asia collision. All four of our main models (M1–M4: Table 2) exhibited similar370
P–T–t paths to one another, as did all but ten of our additional 38 simulations. Key371
deviations occurred when the initial angle of the weak zone was increased, leading to372
tracked continental crustal nodes consistently entering the Forbidden Zone (Fig. 6), which373
is unexpected in modern-day subduction (Liou et al., 2000). Such simulations produced374
maximum pressures >5 GPa for both upper and lower crustal components: however, these375
high weak zone angles are unrealistic for the uppermost portions of modern-day subduction376
systems (Syracuse et al., 2010), and the most deeply subducted continental crust in our377
simulations were small fragments dragged into the mantle during slab breako↵, which378
would never be exhumed. The P–T paths of reasonably exhumable felsic379
rocks—represented here by those that stagnated at the buoyancy-driven threshold for the380
subduction of the upper crust at around 100 km depth (Agard et al., 2009)—exhibited381
peak temperatures of 600–700 �C in all simulations (Fig. 6). Generally shallower382
subduction resulted from applying a ‘mafic granulite’ flow law to the upper and/or lower383
crust (Fig. 6, Table S4), although this is not a representative rheology for the felsic384
lithologies that comprise the bulk of the Tso Morari terrane.385
Considered together, these results indicate that the general shape of the P–T paths386
calculated in simulations M3–M4 (Fig. 2) is robust within the range of reasonable natural387
variation in model parameters. Modifications to our preferred Tso Morari-specific388
geometric and geological constraints (e.g. slab dip angle, convergence velocity) did cause389
15
large variations in calculated P–T–t conditions; however, these alternative paths mostly390
failed to intersect or terminate (reasonably close to) any of the calculated P–T conditions391
proposed for the evolution of the massif, in contrast to the results of our main simulations392
(Fig. 2). Furthermore, these extreme variants are unsupported by independent datasets393
(e.g. U–Pb geochronology), which must provide absolute constraints on our region-specific394
investigation.395
4.2 Implications for the Tso Morari massif396
4.2.1 Orogen-parallel correlation in the northwestern Himalaya397
The geological evolution of the Tso Morari massif is strongly debated, owing to wide398
variation in interpreted depths of subduction and thermal gradients of metamorphism.399
Such geodynamic criteria are important for formulating regional-scale numerical models of400
India–Asia collision, which are commonly benchmarked against P–T–t data reported for401
these Himalayan UHP eclogites (e.g. Warren et al., 2008). In our main models (M3–M4),402
peak eclogite-facies P–T conditions of ⇠2.6–2.8 GPa and ⇠620 �C are reached ⇠3 Myr403
after subduction of continental crust initiated (Fig. 2). This correlates well with ⇠2.4–3.2404
GPa and ⇠650–750 �C reported for eclogites in the UHP Kaghan (Kaneko et al., 2003) and405
HP Stak (Lanari et al., 2013) valleys located 400 km northwest along strike (Fig. 1, inset).406
If pressures up to 4.8 GPa (cf. Mukherjee et al., 2003; Wilke et al., 2015) for peak407
metamorphism at Tso Morari are true, the upper crusts in each locality must have behaved408
independently during subduction. A strong gradient in plate-convergence velocity would409
thus be required between both localities over a relatively small length scale, which we have410
shown produces a significantly di↵erent P–T path (Fig. 6). This seems unreasonable due411
to their current proximity, equivalent orogen-parallel tectonic setting, and similarities in the412
ages and durations of subduction and exhumation. As such, we suggest that both regions413
of the northwest Himalaya represent now-dismembered parts of a single, coherent crustal414
terrane that underwent deep subduction to UHP conditions of ⇠ 2.8 GPa, comparable in415
size to the Western Gneiss Region, Norway, and the Dabie-Sulu terrane, China.416
16
How then can these extreme pressure estimates be reconciled, given that all samples417
studied from the massif must have experienced the same tectonothermal evolution? We418
envisage two possibilities: tectonic overpressure or mineralogical disequilibrium. For the419
former, recent studies have shown that localised, non-lithostatic overpressure may be420
prevalent in some convergent tectonic settings (Petrini and Podladchikov, 2002; Li et al.,421
2010), and small-scale and rigid petrological elements may act as foci for its accumulation422
(Reuber et al., 2016). Importantly, our simulations considered a laterally homogeneous and423
mechanically weak crust, and the maximum modelled tectonic overpressure at the424
resolution of our model never exceeded 0.5 GPa (Fig. 7a–b), so cannot account for the425
discrepant pressure estimates. Although, smaller-scale heterogeneities may result in426
significantly higher localised tectonic pressures (Reuber et al., 2016), such heterogeneities427
in the Tso Morari massif would likely have been on the scale of individual boudins (i.e.428
tens of meters). Considering such features in a lithospheric-scale simulations is429
computationally challenging and was unable to be tested at this point in time, although430
indicates an avenue for future study.431
The latter possibility is mineralogical disequilibrium. We note that both studies432
claiming anomalously high peak pressures for the region (Mukherjee et al., 2003; Wilke et433
al., 2015) applied conventional thermobarometry to minerals that were not conclusively434
proven to have ever been in chemical equilibrium with one another, thus rendering the P–T435
results of questionable veracity. Indeed, final equilibration of Tso Morari crust at 80–100436
km depth, as suggested in this study, produces lithostatic pressures straddling the437
HP–UHP transition (Fig. 2). This result supports numerous petrographic observations438
that some eclogite boudins in the region lack coesite in the peak assemblage, containing439
monocrystalline quartz inclusions in garnet outer rims instead (De Sigoyer et al., 1997;440
St-Onge et al., 2013).441
4.2.2 Metastability during the subduction of felsic crust442
Our petrological modelling suggests that the Polokongka La granite protolith did not443
pervasively transform during subduction, as indicated by the absence (or evidence for the444
17
former existence) of diagnostic high-P phases preserved in the orthogneiss today. This445
non-transformation is attributed to a lack of free fluid, low temperatures associated with446
the early stages of subduction (<400 �C), and the short timescale involved for447
metamorphism (<7 Myr): all of which promote sluggish reaction kinetics (Rubie, 1986).448
By contrast, the Panjal Trap basalt must have pervasively transformed at UHP449
conditions, as shown by the occurrence of coesite in some eclogite boudins. Excess H2
O450
would be liberated from the basalt relatively late in the subduction history at ⇠2.3 GPa451
and ⇠520 �C (Figs 3 and 5a) following the breakdown of chlorite and amphibole. This fluid452
would then have locally infiltrated the enclosing granite and promoted minor mineralogical453
change in the host. As such, fluid-catalysed transformation in the granite may have only454
occurred in thin rinds along boudin–host contacts due to the volumetrically negligible455
proportion of metabasite (<1%) in the terrane as a whole. Any prograde or retrograde456
reactions occurring in the bulk of the granite would have been controlled by its own457
(sluggish) rate of internal dehydration. Jadeitic clinopyroxene, which should have458
comprised ⇠18% of the orthogneiss UHP assemblage under equilibrium conditions (Fig.459
5b), has never been reported from Tso Morari, nor have any indication of its former460
existence, such as relic inclusions or amphibole-bearing symplectites (Epard and Steck,461
2008). Nonetheless, evidence of any fluid-absent transformation that may have occurred462
during subduction would likely have been lost during retrograde low-P amphibolite-facies463
metamorphic overprinting and recrystallisation that occurred under H2
O-saturated and/or464
suprasolidus conditions (Figs 2, 3, and S2: Palin et al., 2014b).465
Similar large-scale metastability of subducted continental crust has been inferred by466
numerous studies of UHP terranes worldwide, including the Western Gneiss Region,467
Norway (Peterman et al., 2009; Young and Kylander-Clark, 2015). The common, rather468
than exceptional, non-transformation of felsic crust during subduction has important469
implications for plate tectonic processes and the geodynamic simulations describing them.470
For example, models of crustal strength based on feldspar-absent assemblages at high-P471
conditions (Stockhert and Renner, 1998), and those invoking fluid-fluxed melting472
(Labrousse et al., 2011), retrogression, or hydrolytic/transformational weakening (Warren,473
18
2013) to instigate exhumation of UHP terranes by a↵ecting the mechanical behaviour and474
rheology of high-P mineral assemblages may not be applicable if vast swathes of475
continental material remain metastably dry, solid, and undeformed at eclogite-facies476
depths. Our petrological modelling suggests that the majority of the Tso Morari terrane477
would have resisted transformation at UHP conditions owing to its fluid-undersaturated478
state, and so is unlikely to have pervasively weakened or strengthened either. Field479
observations in the massif (and from other large-scale continental terranes) provide480
first-hand evidence that internal deformation of UHP sheets is uncommon, with shear481
largely confined to the edges of internally low-strained blocks (e.g. De Sigoyer et al., 2004;482
Epard and Steck, 2008). As such, metamorphic reactions are not simply passive recorders483
of plate tectonic processes, but play an active role in orogenesis; a point that geodynamic484
models involving subduction should consider. Close integration with phase equilibrium485
forward modelling of suitable rock types o↵ers a robust way to achieve this.486
5 CONCLUSIONS487
Integrated petrological and geodynamic forward models of the tectonometamorphic488
evolution of the Himalayan UHP Tso Morari massif predicts peak P–T conditions of489
⇠2.6–2.8 GPa and ⇠600–620 �C. These conditions are consistent with observed490
metamorphic assemblages in mafic eclogite, and were simulated to have been achieved491
within timescales consistent with the reported duration of subduction. Significantly higher492
pressures suggested for peak metamorphism (up to ⇠4.8 GPa) are interpreted to be493
spurious and likely resulted from thermobarometry performed on minerals that were not in494
chemical equilibrium. Our models with homogeneous crustal properties suggest that495
large-scale tectonic overpressures are of insu�cient magnitude to account for this496
discrepancy. Small, boudin-scale variations in mechanical strength—not resolved in our497
simulations—may cause more significant overpressures (Reuber et al., 2016). Yet, the498
proximity, equivalent orogen-parallel tectonic setting, and similar P–T conditions and499
timing of peak metamorphism between Tso Morari and the Kaghan Valley terrane suggest500
that they acted as a coherent, ⇠400-km-long crustal domain that was subducted to UHP501
19
conditions during the earliest stages of India–Asia collision.502
Our petrological modelling shows that degree of fluid saturation of each massif503
protolith imparts a strong control on their transformation potential during subduction.504
The relatively high H2
O content of Panjal Trap basalt—the protolith for the mafic505
eclogite—allowed generation of free fluid and complete mineralogical transformation at506
eclogite-facies conditions. The relatively low H2
O content of Polokongka La granite—the507
protolith of the orthogneiss that now hosts mafic eclogite boudins—would have inhibited508
transformation. It therefore likely persisted as a metastable granite during much of the509
subduction–exhumation cycle. Widespread metamorphic change must have occurred510
during retrograde residence in the crust, as supported by its characteristic511
amphibolite-facies assemblages and lack of evidence for the former existence of diagnostic512
high-pressure phases predicted by petrological modelling (e.g. coesite, jadeitic513
clinopyroxene). The volumetric dominance of dry, felsic lithologies in continental (U)HP514
terranes worldwide indicates that the large-scale metastability of subducted crust may be515
common during collisional orogenesis. Geodynamic simulations of tectonic processes516
predicated on the operation of equilibrium metamorphism should consider the e↵ects of517
evolving fluid budgets in controlling reaction catalysis and the variable transformation of518
rocks at depth in the Earth.519
6 ACKNOWLEDGEMENTS520
Reviews by Stephan Guillot, Julia De Sigoyer, Mary Leech, Clare Warren, and an521
anonymous reviewer are gratefully acknowledged. Michael Bickle is thanked for editorial522
handling. This research did not receive any specific grant from funding agencies in the523
public, commercial, or not-for-profit sectors.524
20
References525
Agard, P., Yamato, P., Jolivet, L., Burov, E., 2009. Exhumation of oceanic blueschists and526
eclogites in subduction zones: Timing and mechanisms. Earth-Science Reviews 92,527
53–79.528
Berthelsen, A., 1953. On the geology of the Rupshu district, NW Himalaya. Meddelelser529
fra Dansk Geologisk Forening 12, 350–415.530
Carswell, D. A., Cuthbert, S. J., 1986. Eclogite-facies metamorphism in the lower531
continental crust. In: Dawson, J. B. et al. (Ed.), The Nature of the Lower Continental532
Crust. Vol. 24. Blackwell Scientific Publications, Oxford, UK, pp. 193–209.533
Chatterjee, N., Jagoutz, O., 2015. Exhumation of the UHP Tso Morari eclogite as a diapir534
rising through the mantle wedge. Contributions to Mineralogy and Petrology 169, 1–20.535
Dal Piaz, G. V., Lombardo, B., 1986. Early-alpine eclogite metamorphism in the Penninic536
Monte Rosa-Gran Paradiso basement nappes of the north-western Alps. GSA Memoirs537
164, 249–266.538
De Sigoyer, J., Chavagnac, V., Blichert-Toft, J., Villa, I. M., Luais, B., Guillot, S., Cosca,539
M., 2000. Dating the Indian continental subduction and collisional thickening in the540
north-west Himalaya: Multichronology of the Tso Morari eclogites. Geology 28, 487–490.541
De Sigoyer, J., Guillot, S., Dick, P., 2004. Exhumation of the ultrahigh-pressure Tso542
Morari unit in eastern Ladakh (NW Himalaya): A case study. Tectonics 23, TC3003.543
De Sigoyer, J., Guillot, S., Lardeaux, J. M., Mascle, G., 1997. Glaucophane-bearing544
eclogites in the Tso Morari dome (eastern Ladakh, NW Himalaya). European Journal of545
Mineralogy 9, 1073–1083.546
Drucker, D. C., Prager, W., 1952. Soil mechanics and plastic analysis for limit design.547
Quarterly of Applied Mathematics 10, 157–165.548
Epard, J. L., Steck, A., 2008. Structural development of the Tso Morari ultra-high pressure549
nappe of the Ladakh Himalaya. Tectonophysics 451, 242–264.550
21
Gerya, T., Stockhert, B., 2006. Two-dimensional numerical modeling of tectonic and551
metamorphic histories at active continental margins. International Journal of Earth552
Sciences 95, 250–274.553
Gerya, T., Stockhert, B., Perchuk, A. L., 2002. Exhumation of high-pressure metamorphic554
rocks in a subduction channel: a numerical simulation. Tectonics 21, 1056.555
Girard, M., Bussy, F., 1999. Late Pan-African magmatism in the Himalaya: new556
geochronological and geochemical data from the Ordovician Tso Morari metagranites557
(Ladakh, NW India). Schweizerische Mineralogische und Petrographische Mitteilungen558
79, 399–417.559
Guillot, S., De Sigoyer, J., Lardeaux, J. M., Mascle, G., 1997. Eclogitic metasediments from560
the Tso Morari area (Ladakh, Himalaya): evidence for continental subduction during561
India–Asia convergence. Contributions to Mineralogy and Petrology 128, 197–212.562
Guillot, S., Garzanti, E., Baratoux, D., Marquer, D., Maheo, G., De Sigoyer, J., 2003.563
Reconstructing the total shortening history of the NW Himalaya. Geochemistry,564
Geophysics, Geosystems 4, 1064.565
Holland, T. J. B., Powell, R., 1998. An internally consistent thermodynamic dataset for566
phases of petrological interest. Journal of Metamorphic Geology 16, 309–343.567
Jagoutz, O., Royden, L., Holt, A. F., Becker, T. W., 2015. Anomalously fast convergence568
of India and Eurasia caused by double subduction. Nature Geoscience 8, 475–478.569
Kaneko, Y., Katayama, I., Yamamoto, H., Misawa, K., Ishikawa, M., Rehman, H. U.,570
Kausar, A. B., Shiraishi, K., 2003. Timing of Himalayan ultrahigh-pressure571
metamorphism: sinking rate and subduction angle of the Indian continental crust572
beneath Asia. Journal of Metamorphic Geology 21, 589–599.573
Kaus, B. J. P., 2010. Factors that control the angle of shear bands in geodynamic574
numerical models of brittle deformation. Tectonophysics 484, 36–47.575
Kirby, S. H., Durham, W. B., Stern, L. A., 1991. Mantle phase changes and576
deep-earthquake faulting in subducting lithosphere. Science 252, 216–225.577
22
Labrousse, L., Prouteau, G., Ganzhorn, A. C., 2011. Continental exhumation triggered by578
partial melting at ultrahigh pressure. Geology 39, 1171–1174.579
Lanari, P. L., Riel, N., Guillot, S., Vidal, O., Schwartz, S., Pecher, A., Hattori, K. H., 2013.580
Deciphering high-pressure metamorphism in collisional context using microprobe581
mapping methods: Application to the Stak eclogitic massif (northwest Himalaya).582
Geology 41, 111–114.583
Leech, M., Singh, S., Jain, A. K., Klemperer, S. L., Manickavasagam, R. M., 2005. The584
onset of India–Asia continental collision: Early, steep subduction required by the timing585
of UHP metamorphism in the western Himalaya. Earth and Planetary Science Letters586
234, 83–97.587
Li, Z. H., Gerya, T. V., and Burg, J. P., 2010. Influence of tectonic overpressure on P–T588
paths of HP–UHP rocks in continental collision zones: thermomechanical modelling.589
Journal of Metamorphic Geology 28, 227–247.590
Liou, J. G., Hacker, B. R., Zhang, R. Y., 2000. Into the Forbidden Zone. Science 287,591
1215–1216.592
Marquardt, H., Miyagi, L., 2015. Slab stagnation in the shallow lower mantle linked to an593
increase in mantle viscosity. Nature Geoscience 8, 311–314.594
Maruyama, S., Liou, J. G., Terabayashi, M., 1996. Blueschists and eclogites of the world595
and their exhumation. International Geology Review 38, 485–594.596
Mascle, G., Colchen, M., Guillot, S., Delaygue, G., 1994. Preliminary results on the597
metamorphic evolution of the Tso Morari dome: Consequences for the geodynamic598
evolution of the Indian margin during Himalayan collision. Nepal Geological Society599
Journal 10, 85–86.600
Mukherjee, B. K., Sachan, H. K., Ogasawara, Y., Muko, A., Yoshioka, N., 2003.601
Carbonate-bearing UHPM rocks from the Tso Morari region, Ladakh, India: petrological602
implications. International Geology Review 45, 49–69.603
23
O’Brien, P. J., Rotzler, J., 2003. High-pressure granulites: formation, recovery of peak604
conditions and implications for tectonics. Journal of Metamorphic Geology 21, 3–20.605
Palin, R. M., St-Onge, M. R., Waters, D. J., Searle, M. P., Dyck, B., 2014. Phase equilibria606
modelling of retrograde amphibole and clinozoisite in mafic eclogite from the Tso Morari607
massif, northwest India: constraining the P–T–M(H2
O) conditions of exhumation.608
Journal of Metamorphic Geology 32, 675–693.609
Palin, R. M., Weller, O. M., Waters, D. J., Dyck, B., 2016. Quantifying geological610
uncertainty in metamorphic phase equilibria modelling; a Monte Carlo assessment and611
implications for tectonic interpretations. Geoscience Frontiers 7, 591–607.612
Peacock, S. M., 1989. Thermal modeling of metamorphic pressure–temperature–time613
paths: A forward approach. In: Spear, F. S., Peacock, S. M. (Eds.), Metamorphic614
Pressure–Temperature–Time Paths. Vol. 7. American Geophysical Union Short Course615
Series, Washington, DC, pp. 57–102.616
Peterman, E. M., Hacker, B. R., Baxter, E. F., 2009. Phase transformations of continental617
crust during subduction and exhumation: Western Gneiss Region, Norway. European618
Journal of Mineralogy 21, 1097–1118.619
Petrini, K., Podladchikov, Y., 2002. Lithospheric pressure–depth relationship in620
compressive regions of thickened crust. Journal of Metamorphic Geology 18, 67–77.621
Powell, R., Holland, T. J. B., 1988. An internally consistent dataset with uncertainties and622
correlations: 3. Applications to geobarometry, worked examples, and a computer623
program. Journal of Metamorphic Geology 6, 173–204.624
Powell, R., Holland, T. J. B., 2008. On thermobarometry. Journal of Metamorphic625
Geology 26, 155–179.626
Press, W. H., Flannery, B. P., Teukolsky, S. A., Vetterling, W. T., 1992. Runge-Kutta627
Method. Numerical Recipes in FORTRAN: The Art of Scientific Computing. Cambridge628
University Press, Ch. 16.1–16.2, pp. 704–716.629
24
Reuber, G., Kaus, B. J. P., Schmalholz, S. M., White, R. W., 2016. Nonlithostatic pressure630
during subduction and collision and the formation of (ultra)high-pressure rocks. Geology631
44, 343–346.632
Rubie, D. C., 1986. The catalysis of mineral reactions by water and restrictions on the633
presence of aqueous fluid during metamorphism. Mineralogical Magazine 50, 399–415.634
Shellnutt, J. G., Bhat, G. M., Wang, K. L., Brookfield, M. E., Jahn, B. M., Dostal, J.,635
2014. Petrogenesis of the flood basalts from the Early Permian Panjal Traps, Kashmir,636
India: Geochemical evidence for shallow melting of the mantle. Lithos 204, 159–171.637
St-Onge, M. R., Rayner, N., Palin, R. M., Searle, M. P., Waters, D. J., 2013. Integrated638
pressure–temperature–time constraints for the Tso Morari dome (NW India):639
Implications for the burial and exhumation path of UHP units in the western Himalaya.640
Journal of Metamorphic Geology 31, 469–504.641
Stockhert, B., Renner, J., 1998. Rheology of crustal rocks at ultrahigh pressure. In: Hacker,642
B. R., Liou, J. G. (Eds.), When Continents Collide: Geodynamics and Geochemistry of643
Ultrahigh-Pressure Rocks. Kluwer Academic Publishers, Dordrecht, pp. 57–95.644
Syracuse, E. M., van Keken, P. E., Abers, G. A., 2010. The global range of subduction645
zone thermal models. Physics of the Earth and Planetary Interiors 183, 73–90.646
Thielmann, M., Kaus, B. J. P., 2012. Shear heating induced lithospheric-scale localization:647
does it result in subduction? Earth and Planetary Science Letters 359–360, 1–13.648
Warren, C. J., 2013. Exhumation of (ultra-)high-pressure terranes: concepts and649
mechanisms. Solid Earth 4, 75–92.650
Warren, C. J., Beaumont, C., Jamieson, R. A., 2008. Modelling tectonic styles and651
ultra-high pressure (UHP) rock exhumation during the transition from oceanic652
subduction to continental collision. Earth and Planetary Science Letters 267, 129–145.653
Whitney, D., Evans, B., 2010. Abbreviations for names of rock-forming minerals. American654
Mineralogist 95, 185–187.655
25
Wilke, F. D. H., O’Brien, P. J., Schmidt, A., Ziemann, M. A., 2015. Subduction, peak and656
multi-stage exhumation metamorphism: Traces from one coesite-bearing eclogite, Tso657
Morari, western Himalaya. Lithos 231, 77–91.658
Young, D. J., Kylander-Clark, A. R. C., 2015. Does continental crust transform during659
eclogite facies metamorphism? Journal of Metamorphic Geology 33, 331–357.660
26
Figures661
Ordovician granite
GHS
Thrust fault
Normal fault
Neotethyan shelfTso Morari
massif (eclogite, orthogneiss,
metasediment)
Mélange and ophiolite
TownRiver
Karakoram Range
Trans-Himalayan batholith
HIG
HH
IMAL
AYA
LehLeh
PadumPadum
lakelake
Ladakh batholith
Ladakh batholith
Zanzkar Range (Tethyan Series)
Zanzkar Range (Tethyan Series)
Shyok suture zone
Shyok suture zone
Indus suture zone
Indus suture zone
77°E77°E
78°E78°E
34°N34°N
77°E77°E 78°E78°E
NN
40 km40 km
Panjal Trapbasalt
Panjal TrapbasaltZanskar
River
ZanskarRiver
Tso Morari
Tso Morari
30°N
KG ST
INDIA
TIBET30°N
75°E 80°E
HIMALAYAN RANGEMFT
ISZ
PolokongkaLa
PolokongkaLa
lakelake
Figure 1: Simplified geological map of the Tso Morari region, northwest India, modified after
Epard and Steck (2008), and Shellnutt et al. (2014). Inset shows its location within the Himalayan–
Tibetan orogen. GHS = Greater Himalayan Series, ISZ = Indus Suture Zone, KG = Kaghan Valley,
MFT = Main Frontal Thrust, ST = Stak Valley.
27
Temperature (°C)600 800700100 300 400 500 1000900200
Pres
sure
(GPa
)
3.5
0.5
1.0
1.5
2.0
2.5
3.0
0
5.0
4.0
4.5
Approximate depth assum
ing lithostatic pressure (km)
48
112
32
16
64
144
160
128
0
96
80
150 °
C/GPa
Jd + Qtz
Ab
CoeCoe
Qtz
GrDia
Syra
cuse et al. (2010)
Zo-EclZo-EclLawso
niteec
logite
Lawso
niteec
logite
Dry eclogiteDry eclogite
Amp-EclAmp-Ecl
GranuliteGranuliteAmphiboliteAmphibolite
Ep-Amp
Ep-Amp
GsGs
PPPPZeoZeo
Diamondeclogite
Diamondeclogite
UH
TU
HT
Blueschist
Blueschist
High-Pgranulite
High-Pgranulite
W15
M03
S13
G97D97
G97 D97S13 W15
CJ15
0 Myr0 Myr
2 Myr2 Myr
3–7 Myr3–7 Myr
1
2
3
0
200 400 600
T (°C)
P (G
Pa)
QtzQtzCoeCoe
1 Myr1 Myr
CJ15
Full t
ransfo
rmat
ion
Full t
ransfo
rmat
ion
Non-transf
orm
ation
Non-transf
orm
ation
Full t
rans
form
atio
n
Full t
rans
form
atio
n
Non-tran
s
Non-tran
s
Figure 2: Pressure–temperature (P–T ) plot of reported metamorphic conditions for Tso Morari
units alongside our thermomechanically modelled prograde P–T paths M3 (non-transformation)
and M4 (full transformation). Inset shows elapsed simulation time at 1-Myr intervals along each
path. Dashed black arrow continuing to higher pressures represents a non-calculated theoretical
extension of these paths. Thick dashed white arrow represents an estimated exhumation path.
Boxes indicate estimated P–T conditions of metamorphism during subduction and exhumation:
D97 = De Sigoyer et al. (1997), G97 = Guillot et al. (1997), M03 = Mukherjee et al. (2003), S13
= St-Onge et al. (2013), CJ15 = Chatterjee and Jagoutz (2015), and W15 = Wilke et al. (2015).
Facies grid is modified after Maruyama et al. (1996): Zeo = zeolite, PP = prehnite–pumpellyite,
Gs = greenschist, Ep-Amp = epidote-amphibolite, Amp-Ecl = amphibole-eclogite, and Zo-Ecl =
zoisite-eclogite. Band labeled Syracuse et al. (2010) represents P–T profiles for descending slab
surfaces in modern-day subduction zones. Geotherm marked 150 �C/GPa denotes the limit of the
Forbidden Zone (Liou et al., 2000).
28
2010 0000
3.00
2.80
2.60
2.952.95
2.652.65
3.05
CoeCoeQtzQtz
70
0
70
60609090
202010103.503.50
3.40
3.153.15
3.553.55
3.35
3.503.50
CoeCoeQtzQtz
3.003.002.952.95
Temperature (°C)500 600 800700300 400
3.7
2.7
2.9
3.1
3.3
3.5
2.5
2.8
3.0
3.2
3.4
3.6
2.6Bulk
-rock
den
sity
(0.0
01 ×
kg/
m3 )
100
0
>0
10
20
30
40
50
60
70
80
90
Bulk
-rock
H2O
as a
free
pha
se (%
)
1.0
2.0
3.0
0.2
5.0
4.0
Pres
sure
(GPa
)
POLOKONGKA LA GRANITE PANJAL TRAP BASALT
1.0
2.0
3.0
5.0
4.0
Pres
sure
(GPa
)
Temperature (°C)500 600 800700300 400
Figure 3: Variations in bulk-rock water contents and densities for metamorphosed Polokongka
La granite and Panjal Trap basalt protoliths. Fully labeled phase diagrams are presented in Figs
S1–4. Stippled regions represent the global range of P–T profiles for descending slab surfaces
in modern-day subduction zones (Syracuse et al., 2010). Thick, dashed, white arrows represent
an estimated exhumation path reproduced from De Sigoyer et al. (2000). Grey boxes represent
previously reported P–T conditions of peak metamorphism, as outlined in Fig. 2.
29
0 4000Width (km)
No stress
Free slip
Free
slip
Free slip
Dep
th (k
m)
0
200
400
660
Upper crustLower crust
Lithospheric mantle
Asthenospheric mantle
Weak channel
100 °C
500 °C
900 °C
1300 °C
1700 1800 1900 2000 2100
Figure 4: Initial setup for geodynamic modelling. Lithospheric convergence was controlled by a
pushing force directed from right-to-left (red box and arrow). A weak channel was included in order
to initialise subduction. 1 and 2 = upper crust, 3 and 4 = lower crust, 5 = lithospheric mantle, 6 =
aesthenospheric mantle, and 7 = weak channel. Detailed information about rheological parameters
is given in Tables S2 and S3 in the Supplementary Information.
30
Temperature (°C)470 520 570370 420 620320270 670 720 770
Mod
al p
ropo
rtio
n
0.4
0.6
0.8
1.0
0.2
0
Density (ρ) (0.001 × kg/m
3)
3.3
3.4
3.8
3.2
3.1
3.5
3.6
3.7
2.42.0 2.82.61.4 1.6 1.8 4.83.2 3.6 4.0 4.4Pressure (GPa)
sub-
met
amor
phic
/rap
idly
dec
reas
ing
relia
bilit
ysu
b-m
etam
orph
ic/r
apid
ly d
ecre
asin
g re
liabi
lity
A
Grt rim
Ctd
GlnAct
Omp
Tlc
Ms
BtChl
ChlEp
Rt H2OSpnLws
Ms
Coe
Brs–Wnc
H2O
Grt rimGrt core Grt core
ρbulk-rockFR
AC
sub-
met
amor
phic
/rap
idly
dec
reas
ing
relia
bilit
ysu
b-m
etam
orph
ic/r
apid
ly d
ecre
asin
g re
liabi
lity
Mod
al p
ropo
rtio
n
0.4
0.6
0.8
1.0
0.2
0
Temperature (°C)470 520 570370 420 620320270 670 720 770
Density (ρ) (0.001 × kg/m
3)
2.9
3.0
3.1
2.8
2.6
2.7
2.42.0 2.8Pressure (GPa) 2.61.4 1.6 1.8 4.83.2 3.6 4.0 4.4B
Omp
CoeQtz
Ms
Kfs
Grt
Jd
Ky
Rt
Jd
Ms
Kfs
Ky
Grt
ρbulk-rock
Figure 5: Changes in calculated mineral proportions for both protoliths during metamorphism
along the calculated M4 simulation P–T path (up to ⇠2.8 GPa and ⇠620 �C) and its theoretical
extension up to 4.8 GPa and ⇠770 �C shown in Fig. 2. A: Panjal Trap basalt. B: Polokongka
La granite. Bulk-rock densities (dashed lines labeled ⇢bulk-rock
) assume complete transformations.
Mineral abbreviations are after Whitney and Evans (2010): Act, actinolite; Brs, barroisite; Bt,
biotite; Chl, chlorite; Coe, coesite; Ctd, chloritoid; Ep, epidote; Gln, glaucophane; Grt, garnet;
Jd, jadeite; Kfs, K-feldspar; Ky, kyanite; Lws, lawsonite; Ms, muscovite; Omp, omphacite; Qtz,
quartz; Rt, rutile; Spn, sphene; Tlc, talc; Wnc, winchite. FRAC denotes the P–T conditions at
which garnet cores were fractionated out of the eclogite bulk composition.
31
Temperature (°C)
600 800700100 300 400 500 9002000
Pres
sure
(GPa
)
1.0
2.0
3.0
0
5.0
4.0
6.0
Mean P–T path for entiresubducted upper crustin each simulation
Jd + Qtz
150 °C/G
Pa
Qtz
DiaGr
Coe
Ab
Forbidden zoneSy
racu
se et
al. (2
010)
High weakzone angle
Temperature-dependent conductivity
Hydratedcrust
Temperature (°C)
600 800700100 300 400 500 9002000
Pres
sure
(GPa
)
1.0
2.0
3.0
0
5.0
4.0
Jd + Qtz
150 °
C/GPa
Ab
P–T paths for reasonablyexhumable upper-crustal rocks in eachsimulation
Syra
cuse
et al
. (2010)
High platevelocity
High weakzone angle
Mafic granulitecrustal flow law
Temperature-dependent conductivity
T-dependentconductivity
GrDia
Temperature (°C)
200 600 800 1000 12004000
Pres
sure
(GPa
)
2.0
3.0
4.0
5.0
6.0
0
10.0
7.0
1.0
8.0
9.0
Jd + Qtz
150 °C/G
Pa
Coe
DiaGr
Qtz
Ab
P–T path for highest-pressure upper-crustalnode in each simulation
Syra
cuse
et a
l. (2
010)
High weakzone angle
Forbiddenzone
Temperature-dependent
conductivity
Temperature (°C)
200 600 800 1000 12004000
Pres
sure
(GPa
)
2.0
3.0
4.0
5.0
6.0
0
8.0
7.0
1.0
Jd + Qtz
150 °C/G
Pa
Coe
DiaGr
Qtz
Ab
Forbiddenzone
P–T path for highest-temperature upper-crustalnode in each simulation
Syra
cuse
et
al. (
2010
)
High weakzone angle
Temperature-dependent
conductivity
Figure 6: Results of numerical modelling sensitivity tests. Each arrow represents a P–T path
calculated for a di↵erent set of model parameters. Red arrows represent P–T paths calculated for
simulations M1–M4. Band labelled Syracuse et al. (2010) represents the range of P–T profiles for
descending slab surfaces in modern-day subduction zones, and gray boxes show P–T metamorphic
conditions reported for the Tso Morari massif by various studies (cf. Fig. 2). Top left: Mean
P–T paths for the whole subducted upper crust. Blue squares represent one standard deviation
in temperature and pressure. Top right: P–T paths of reasonably exhumable rocks. Bottom row:
P–T paths for nodes that reached the maximum pressure within the subducted upper crust (left)
and lower crust (right) in each simulation.
32
CRUST
coesitequartz
quartz
coesite
CRUST
MANTLE
MOHO
Maximum overpressure (GPa)
0 0.2 0.4 0.6 0.8 1.0−1.0 −0.8 −0.6 −0.4 −0.2
1650 1700 1750 1800 1850 1900 1950 2000 2050 2100 2150 2200
Dep
th (k
m)
0
50
100
150
200
7 Myr
ASIAN PLATE INDIAN PLATE
SUTUREZONE
TSO MORARIMASSIF UNITS
Width (km)1650 1700 1750 1800 1850 1900 1950 2000 2050 2100 2150 2200
MANTLE
Dep
th (k
m)
0
50
100
150
200
1.2−1.2
A
BTSO MORARI
MASSIF UNITS
Figure 7: Results of integrated petrological–geodynamic modelling for simulation M3. A: Geo-
dynamic model components after 7 Myr of subduction. Green square represents the locus of the
Tso Morari massif units under simulation M3 conditions after this time period, and blue hexagon,
shown for reference, represents these units for simulation M4. B: Maximum overpressure recorded
by each individual node. Solid black lines demarcate the crust and dashed black line represents
a P–T locus for the quartz–coesite transition. Note that the lower crust in these simulations is
mechanically homogeneous and internal small-scale heterogeneities are ignored, which may result
in additional overpressures (Reuber et al., 2016).
33
Tables662
Table 1: Bulk-rock compositions used in phase diagram construction for the Polokongka La granite
(Girard and Bussy, 1999) and the Panjal Trap basalt (Shellnutt et al., 2014) given as molar pro-
portions of oxides: aunfractionated composition (no sequestration of elements into garnet cores),
bfractionated bulk composition (following sequestration of elements into garnet cores). FeOtot =
all iron as FeO, such that mol.% O = mol.% Fe2
O3
.
Lithology H2
O SiO2
Al2
O3
CaO MgO FeOtotal K2
O Na2
O TiO2
O
Polokongka La granite 2.04 80.00 8.35 0.68 0.48 1.55 3.88 2.46 0.15 0.40
Panjal Trap basalta 13.11 44.43 9.36 8.74 10.42 8.60 0.36 3.06 1.49 0.43
Panjal Trap basaltb 13.95 44.54 9.07 8.01 11.01 7.76 0.38 3.25 1.59 0.44
Table 2: Summary characteristics of the reference geodynamic models presented in this work.
aprofile after Leech et al. (2005), bprofile modified from Guillot et al. (2003).
Model no. Convergence velocity Phase transformation
M1 Constant (7 cm/yr)a No: constant terrane density of 2700 kg/m3
M2 Constant (7 cm/yr)a Yes: P–T -dependent terrane density (cf. Fig. 3)
M3 Deceleration (10 cm/yr for first No: constant terrane density of 2700 kg/m3
2 Myr and 4 cm/yr afterwards)b
M4 Deceleration (10 cm/yr for first Yes: P–T -dependent terrane density (cf. Fig. 3)
2 Myr and 4 cm/yr afterwards)b
34
1
SUPPLEMENTARY INFORMATION for:
Subduction metamorphism in the Himalayan ultrahigh-pressure Tso
Morari massif: an integrated geodynamic and petrological modelling
approach
Richard M. Palin*1,2, Georg Reuber1, Richard W. White1, Boris J.P. Kaus1 and Owen M.
Weller3 1Institute of Geosciences, Johannes-Gutenberg University of Mainz, 55128 Mainz, Germany 2Department of Geology and Geological Engineering, Colorado School of Mines, Golden,
80401, Colorado, USA 3Department of Earth Sciences, University of Cambridge, Cambridge, CB2 3EQ, UK
*corresponding author: [email protected]
Contents
• Petrological modelling parameters, methodology, and bulk compositions
• Tables S1–S5
• Figures S1–S8
• References
BULK COMPOSITIONS AND ACTIVITY–COMPOSITION RELATIONS USED FOR
METAMORPHIC PHASE EQUILIBRIA MODELLING
Petrological phase diagram modeling utilized the following activity–composition relations for
solid-solution phases: glaucophane, actinolite, hornblende, gedrite, diopside, omphacite, and
jadeite (Diener and Powell, 2012), muscovite and paragonite (Coggon and Holland, 2002), talc
and epidote (Holland and Powell, 1998), chlorite (Holland et al., 1998), cordierite, biotite and
garnet (White et al., 2007), plagioclase and K-feldspar (Holland and Powell, 2003), ilmenite
and hematite (White et al., 2000). Pure phases comprised albite, zoisite, lawsonite, rutile,
sphene, quartz, sillimanite, andalusite, kyanite, and fluid (H2O).
2
Whereas the whole-rock composition for Polokongka La granite sample AS9660 reported by
Girard and Bussy (1999) specified both FeO and Fe2O3 contents, that reported for Panjal Trap
basalt sample PJ2-010 by Shellnutt et al. (2014) presented all iron as Fe2O3. As such, a bulk-
rock XFe3+ = Fe3+(Fe2++ Fe3+) = 2 × O/FeOtotal ratio of 0.1 was applied for our modeling (Table
1 in the main manuscript), in keeping with the observation that unaltered mafic igneous rocks
generally have a low oxidation state (Schilling et al., 1983). The bulk-rock XFe3+ ratio for
Polokongka La granite sample AS9660 was not modified from its reported value.
Finally, fractionation of the eclogite bulk-rock composition owing to sequestration of cations
into growing garnet porphyroblasts during prograde metamorphism was considered in order to
examine the effect of an evolving bulk composition on the sensitivity of calculated phase
equilibria. Distinct compositional breaks across core–rim boundaries in garnet (e.g. St-Onge et
al., 2013; Wilke et al., 2015) were chosen as a key datum point for this procedure. Core
domains were assumed to represent ~20 vol.% of entire grains based on examination of
compositional line profiles in samples of fresh eclogite and assuming a spherical porphyroblast
geometry. Garnet constitutes ~30 vol.% of peak UHP eclogite-facies parageneses (Wilke et al.,
2015), and so a proportion of ~6 vol.% was used to represent the criterion at which core growth
was superceded by rim growth during subduction. Sample PJ2-010 was calculated to form 6%
garnet along the prograde P–T path calculated via geodynamic modelling at ~18.75 kbar (Figs
S3 and S4), and fractionation was performed at this point.
TABLE S1
Lithology SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O LOIe
Polokongka La granitea
74.46 0.19 13.19 1.00 0.83 0.02 0.30 0.59 2.36 5.66 0.57
Panjal Trap basaltb 45.78 2.04 16.37 11.78c N.R.d 0.16 7.20 8.41 3.25 0.58 4.05
Protolith bulk-rock compositions (wt% oxide). aGirard and Bussy (1999), sample AS9660, bShellnutt et al. (2014), sample PJ2-010, cgiven in original study as total iron, dN.R. = not
reported, eLOI = loss on ignition.
3
TABLE S2
Component C (MPa)a Φb Cp (J/K)c k [W/(mK)]d Q (W/m³)e ρ0 (kg/m3)f
Upper crust 1 5.00 1000 2.339 1.750 × 10−6 2700g
Lower crust 1 5.00 1000 1.97 0.250 × 10−6 2900
Lithosphere 1 28.68 1000 1.99 0.022 × 10−6 3300
Asthenosphere 1 28.68 1000 1.99 0.022 × 10−6 3300
Weak channel 1 3.44 1000 2.00 0 3300
Material properties used for the main (reference) numerical simulation. Overall rheological
parameters are assembled from Turcotte and Schubert (2014), Li et al. (2010), and Bittner and
Schmeling (1995). aCohesion, bfriction angle, cheat capacity, dthermal conductivity, eradioactive heat generation, freference density (ρ = f(ρ0, T)), gdensities varied in the case of a
transformation simulation.
TABLE S3
Component Flow lawa A (MPa−n s−1)b nc E (kJ/mol)d V (m³/mol)e
Upper crust Quartzitef 6.7 × 10−6 2.4 156 0
Lower crust Plagioclaseg 3.3 × 10−4 3.2 238 0
Lithosphere Dry olivineh 2.5 × 104 3.5 532 17 × 10−6
Asthenosphere Dry olivineh 2.5 × 104 3.5 532 17 × 10−6
Weak channel Wet olivinei 2.0 × 103 4.0 471 0
Rheological parameters used for the main (standard-state) numerical simulation. aDislocation
creep flow law, bviscosity prefactor, cpower law exponent, dactivation energy, eactivation
volume, fRanalli and Murphy (1987), gShelton and Tullis (1981), hKirby (1983), iRanalli
(1995).
4
TABLE S4
Phas
e tra
nsiti
on
Con
verg
ence
rate
Dec
eler
atio
n
Flow
law
(U
C)
Flow
law
(L
C)
Wea
k zo
ne
angl
e
Eros
ion
rate
Sedi
men
tat
ion
rate
Fric
tion
angl
e (U
C)
Fric
tion
angl
e (L
C)
Q
Cp
Mea
n P
Mea
n T
Peak
P
Peak
T
Comments
(cm/a) (°) (mm/a) (mm/a) (W/m3) (J/K) (GPa) (°C) (GPa) (°C)
Ref_7 n 7 n Quartzite Plagioclase (An75) 15 5 1 5 5 v 1000 0.7 157 5 1030 M1
Ref_7_P y * * * * * * * * * * * 0.6 154 5.1 913 M2 Ref_10 * 10 y * * * * * * * * * 0.6 170 4 623 M3 Ref_10_P y 10 y * * * * * * * * * 0.7 165 4.2 832 M4
Ref_7_Hyd * * * * * * * * * * * * 0.4 120 0.8 220 Similar to reference; hydrated rock
Ref_10_Hyd * 10 * * * * * * * * * * 0.4 135 0.9 240 Similar to reference; hydrated rock
Ref_Cr_5 * 5 * * * * * * * * * * 0.5 154 4 690 Similar to reference Ref_Cr_5_P y 5 * * * * * * * * * * 0.6 152 4 773 Similar to reference Ref_Cr_10 * 10 * * * * * * * * * * 0.9 185 7 1020 Faster subduction Ref_Cr_10_P y 10 * * * * * * * * * * 0.9 178 7.1 989 Faster subduction
Ref_Cr_15 * 15 * * * * * * * * * * 1 211 9.5 1080 Faster subduction, thus higher P
Ref_Cr_15_P y 15 * * * * * * * * * * 1 188 8.2 1080 Faster subduction, thus higher P
Ref_UC_G * * * Mafic granulite * * * * * * * * 0.7 174 4.4 794 Early underthrusting
Ref_UC_G_P y * * Mafic granulite * * * * * * * * 0.7 177 4.6 846 Early underthrusting
Ref_LC_G * * * * Mafic granulite * * * * * * * 0.6 171 3.6 672 Shallow subduction
Ref_LC_G_P y * * * Mafic granulite * * * * * * * 0.6 168 4 764 Shallow subduction
Ref_WA_20 * * * * * 20 * * * * * * 0.7 177 4.4 655 Steep subduction Ref_WA_20_P y * * * * 20 * * * * * * 0.8 176 6.3 1030 Steep subduction
5
Phas
e tra
nsiti
on
Con
verg
ence
rate
Dec
eler
atio
n Fl
ow la
w
(UC
)
Flow
law
(L
C)
Wea
k zo
ne
angl
e
Eros
ion
rate
Sedi
men
tat
ion
rate
Fric
tion
angl
e (U
C)
Fric
tion
angl
e (L
C)
Q
Cp
Mea
n P
Mea
n T
Peak
P
Peak
T
Comments
(cm/a) (°) (mm/a) (mm/a) (W/m3) (J/K) (GPa) (°C) (GPa) (°C) Ref_WA_30 * * * * * 30 * * * * * * 0.6 176 4.2 800 Steep subduction Ref_WA_30_P y * * * * 30 * * * * * * 0.7 173 7 965 Steep subduction
Ref_WA_40 * * * * * 40 * * * * * * 0.7 189 4.4 936 Massively steep subduction
Ref_WA_40_P y * * * * 40 * * * * * * 0.9 198 9.7 1080 Massively steep subduction
Ref_SE_11 * * * * * * 1 * * * * * 0.7 177 4 600 Similar to reference Ref_SE_11_P y * * * * * 1 * * * * * 0.7 170 4.1 601 Similar to reference Ref_SE_15 * * * * * * 1 5 * * * * 0.8 151 4.6 730 Shallow subduction Ref_SE_15_P y * * * * * 1 5 * * * * 0.8 141 4.8 824 Shallow subduction Ref_fa_UC * * * * * * * * 15 * * * 0.6 178 2.3 540 Similar to reference Ref_fa_UC_P y * * * * * * * 15 * * * 0.5 167 2.2 513 Similar to reference Ref_fa_LC * * * * * * * * * 15 * * 0.7 165 4.2 667 Similar to reference Ref_fa_LC_P y * * * * * * * * 15 * * 0.7 160 4.3 628 Similar to reference Ref_Q2 * * * * * * * * * * 2 * 0.5 152 3.1 620 Shallow subduction Ref_Q2_P y * * * * * * * * * 2 * 0.5 148 3.2 621 Shallow subduction Ref_Q0 * * * * * * * * * * 0 * 0.7 142 4.3 650 Similar to reference Ref_Q0_P y * * * * * * * * * 0 * 0.7 140 4.4 860 Similar to reference Ref_CP_1 * * * * * * * * * * * 750 0.7 186 4 684 Similar to reference Ref_CP_1_P y * * * * * * * * * * 750 0.7 182 5.4 987 Similar to reference Ref_CP_2 * * * * * * * * * * * 1250 0.5 136 2.7 584 Shallow subduction Ref_CP_2_P y * * * * * * * * * * 1250 0.5 133 2.7 573 Shallow subduction Ref_7_kT * * * * * * * * * * * * 0.6 248 6.6 1070 Similar to reference
Ref_7_kT_P y * * * * * * * * * * * 0.6 236 6.8 1063 Similar to reference
Ref_10_kT * 10 y * * * * * * * * * 0.7 242 6.2 976 Similar to reference
Ref_10_kT_P y 10 y * * * * * * * * * 0.7 232 6.4 1030 Similar to reference
6
Results of sensitivity testing. Reading from left to right: Phase transition: ‘y’ if phase transitions are applied, ‘n’ if not. Convergence rate (cm/a):
background velocity. Deceleration: ‘y’ if the background velocity is varied with time, ‘n’ if not. Flow law (UC): rheological flow law for the upper
crust. Flow law (LC): rheological flow law for the lower crust. Weak zone angle (°): angle of the weak zone. Erosion rate (mm/a): erosion rate.
Sedimentation rate (mm/a): sedimentation rate. Friction angle (UC): applied friction angle for the upper crust. Friction angle (LC): applied friction
angle for the lower crust. Q (W/m3): heat production for all phases, ‘v’ if it is variable. Cp (J/K): Heat capacity of all phases. Mean pressure: mean
pressure achieved by the subducted upper crust. Mean temperature: Mean temperature achieved by the subducted upper crust. Peak pressure: highest
traced pressure achieved by the subducted upper crust. Peak temperature: highest traced temperature achieved by the subducted upper crust.
Comments: Comments on the evolution of the simulation. All rheological flow laws are taken from Kirby (1983) and Ranalli (1995). The
nomenclature for the simulations is written in the format: “Ref” “Modified parameter”, alongside “P” if phase transitions were applied.
Abbreviations: *: same value as the reference simulation, UC: upper crust, LC: lower crust, P: pressure, T: temperature. Simulations Ref_7,
Ref_7_P, Ref_10, and Ref_10_P (red text) provided the P–T paths discussed in the paper.
TABLE S5
Component k1 k2
Sediments 0.64 807
Upper crust 0.64 807
Lower crust 1.18 474
Lithosphere 0.73 1293
Asthenosphere 0.73 1293
Weak channel 0.73 1293
Parameter values used in simulations that considered temperature-
dependent conductivity (Ref_7_kT, Ref_7_kT_P, Ref_10_kT, and
Ref_10_kT_P: Table S4), which used a conductivity–temperature
relationship k = k1 + k2/(T + 77) from Clauser and Huenges (1995),
where k1 and k2 are empirically derived prefactors and T is the
temperature in Kelvin. These simulations also used the nonlinear
crustal geotherm of McKenzie et al. (2005), which is shown in Fig. S8
7
FIGURE S1
Results of metamorphic mineral equilibria modeling for the bulk-rock composition of Polokongka La granite sample AS9660 (Table 1 in main manuscript text) from Girard and Bussy (1999). (a) Pressure–temperature (P–T) pseudosection. Some small, minor fields are unlabeled for clarity. Assemblage field boundaries marking the appearance or disappearance of magnetite, sphene, ilmenite, hematite, or rutile, are shown as grey dashed lines, which are omitted from the labeled assemblages due to their negligible calculated modal proportions (<1 mol.%). Inset at the top left corner of the diagram shows the P–T relations between these phases. Dotted region represents the global range of descending slab-surface P–T profiles in modern-day subduction zones (Syracuse et al., 2010), and dashed line labeled 150 °C/GPa represents the limit of the Forbidden Zone (Liou et al., 2000). Numbered fields are as follows: 1 – Grt Ms Omp Jd Gln Kfs Ky Qtz, 2 – Grt Ms Omp Jd Kfs Ky Ab Qtz, 3 – Bt Ms Jd Ep Kfs Ab Qtz, 4 – Grt Bt Ms Jd Ep Kfs Ab Qtz, 5 – Bt Ms Pl Kfs Ep Qtz H2O, 6 – Bt Ms Pl Kfs Ab Qtz H2O, 7 – Bt Ms Jd Pl Kfs Qtz, 8 – Grt Ms Jd Pl Kfs Ab Qtz, 9 – Liq Bt Ms Pl Kfs Qtz H2O, 10 – Grt Ms Jd Pl Kfs Qtz, 11 – Liq Grt Bt Ms Pl Kfs Qtz, 12 – Liq Bt Pl Kfs Ky Qtz, 13 – Ms Jd Kfs Ky Qtz.
8
FIGURE S2
Overlay of metamorphic P–T conditions reported for units of the Tso Morari massif (gray and white boxes, and white dashed arrow) and prograde P–T paths calculated via geodynamic modeling (green and yellow lines, and black dashed arrow) onto phase diagrams calculated for Polokongka La granite sample AS9660 (Fig. S1). Acronyms for studies are the same as those listed in the caption to Fig. 2 in the main manuscript. Blue P–T path ending with a hexagon represents that calculated for a time-dependent decreasing subduction velocity (‘deceleration’) and full mineralogical transformation, and green P–T path ending with a square represents that calculated for a time-dependent decreasing subduction velocity (‘deceleration’) with no mineralogical transformation. See the “Geodynamic numerical modeling” section of the main manuscript for more information. All other labels are equivalent to Fig. S1.
9
FIGURE S3
Pressure–temperature (P–T) pseudosection calculated for Panjal Trap basalt sample PJ2-010 from Shellnutt et al. (2014), considering the effects of cation sequestration into garnet cores (Table 1 in the main manuscript text). Some small, minor fields are unlabeled for clarity. Line colors and shadings as for Fig. S1. Bold red line marks the limit of garnet-bearing assemblage fields, and sub-parallel red dashed line marks where garnet comprises 6% of the calculated assemblage using an unfractionated bulk composition. As such, assemblage fields occurring to the high-P side of this contour represent equilibria calculated following fractionation of the effective bulk composition due to removal of garnet core domains. Numbered fields are as follows: 1 – Grt Bt Chl Ctd Tlc Omp Gln Lws, 2 – Grt Bt Chl Ctd Omp Act Gln Lws, 3 – Bt Chl Ep Ctd Omp Act Gln Lws, 4 – Bt Chl Ctd Omp Act Gln Lws, 5 – Chl Ms Ep Omp Gln Act Lws, 6 – Chl Ms Ep Omp Gln Act Lws Qtz, 7 – Bt Chl Ms Ep Omp Qtz Ab, 8 – Bt Chl Ep Act Gln Qtz Ab, 9 – Bt Chl Ep Act Hbl Qtz Ab, 10 – Bt Chl Ms Ep Gln Qtz, 11 – Chl Ms Ep Omp Gln Qtz, 12 – Bt Chl Ms Ep Gln Qtz Zo, 13 – Bt Chl Hbl Pl Qtz, 14 – Bt Chl Act Hbl Pl, 15 – Bt Hbl Kfs Pl H2O, 16 – Bt Opx Hbl Pl H2O, 17 – Bt Ms Hbl Qtz Zo H2O, 18 – Grt Bt Di Hbl Pl Qtz Zo H2O, 19 – Grt Ms Omp Hbl Qtz Zo H2O, 20 – Grt Ms Pg Omp Hbl Qtz Zo H2O, 21 – Chl Ms Ep Omp Gln Qtz Zo, 22 – Grt Chl Ms Ep Omp Gln Act Lws, 23 – Grt Chl Ms Omp Act Gln Lws, 24 – Grt Chl Ms Omp Gln Lws H2O, 25 – Grt Chl Ms Omp Gln Zo H2O, 26 – Grt Ms Omp Gln Zo H2O, 27 – Grt Ms Omp Gln Ky H2O, 28 – Grt Ms
10
Omp Hbl Ky H2O, 29 – Grt Ms Omp Coe Lws Ky H2O, 30 – Grt Ms Omp Coe Lws Tlc H2O, 31 – Grt Ms Tlc Omp Gln Lws H2O, 32 – Grt Bt Ms Tlc Omp Gln Lws H2O, 33 – Grt Bt Chl Tlc Omp Gln Lws H2O.
FIGURE S4
Overlay of metamorphic P–T conditions reported for units of the Tso Morari massif (gray and white boxes, and white dashed arrow) and prograde P–T paths calculated via geodynamic modeling (green and yellow lines, and black dashed arrow) onto phase diagrams calculated for Panjal Trap basalt sample PJ2-010 (Fig. S2). Acronyms for studies are the same as those listed in the caption to Fig. 2 in the main manuscript. Blue P–T path ending with a hexagon represents that calculated for a time-dependent decreasing subduction velocity (‘deceleration’) and full mineralogical transformation, and green P–T path ending with a square represents that calculated for a time-dependent decreasing subduction velocity (‘deceleration’) with no mineralogical transformation. See the “Geodynamic numerical modeling” section of the main manuscript text for more information. All other labels are equivalent to Fig. S2.
11
FIGURE S5
Results of geodynamic numerical modeling for the four end-member scenarios discussed in the
main text: A: Constant velocity with non-transformation (M1), B: Constant velocity with full
transformation (M2), C: Deceleration with non-transformation (M3), D: Deceleration with full
transformation (M4). Insets for each figure part show the modeled evolution of P–T conditions at 1-
Myr time intervals.
12
FIGURE S6
Results of numerical modelling sensitivity tests, color-coded for simulation time (Myr). The data
shown are identical to those in Fig. 6 of the main manuscript, with additional 1-Myr-interval
shading. Note that the upper crustal nodes that reached the highest pressures in each simulation
(bottom left panel) were not necessarily the same as those nodes that reached the highest
temperatures (bottom right panel).
13
FIGURE S7
Evolution of the four simulations at an approximately equal stage of subduction. The upper crust did not subduct to greater depths in any of the
cases. The numerical times of the simulations are highlighted in white within each subfigure. Markers representing the Tso Morari massif units have
the same shape and color as they appear in the main paper and elsewhere in the Supplementary Information.
14
FIGURE S8
Geotherms utilized in geodynamic simulations. Red line represents the non-linear geotherm of
McKenzie et al. (2005), which was utilized in simulations that considered temperature-dependent
conductivity. Blue line represents the constant 9 ºC/km geotherm employed for all other models,
including the reference (preferred) simulations.
15
REFERENCES CITED
Bittner, D., and Schmeling, H., 1995. Numerical modelling of melting processes and induced
diapirism in the lower crust. Geophysical Journal International, 123, 59–70.
Clauser, C., and Huenges, E., 1995. Thermal conductivity of rocks and minerals. Rock physics &
phase relations: A handbook of physical constants, 105–126.
Coggon, R., and Holland, T.J.B., 2002. Mixing properties of phengitic micas and revised garnet-
phengite thermobarometers. Journal of Metamorphic Geology, 20, 683–696.
Diener, J.F.A., and Powell, R., 2012. Revised activity–composition models for clinopyroxene and
amphibole. Journal of Metamorphic Geology, 30, 131–142.
Girard, M., and Bussy, F., 1999. Late Pan-African magmatism in the Himalaya: new
geochronological and geochemical data from the Ordovician Tso Morari metagranites
(Ladakh, NW India). Schweizerische Mineralogische und Petrographische Mitteilungen, 79,
399–417.
Holland, T.J.B., and Powell, R., 1998. An internally-consistent thermodynamic dataset for phases of
petrological interest. Journal of Metamorphic Geology, 16, 309–344.
Holland, T.J.B., and Powell, R., 2003. Activity–composition relations for phases in petrological
calculations: an asymmetric multicomponent formulation. Contributions to Mineralogy and
Petrology, 145, 492–501.
Holland, T.J.B., Baker, J.M., and Powell, R., 1998. Mixing properties and activity–composition
relationships of chlorites in the system MgO–FeO–Al2O3–SiO2–H2O: European Journal of
Mineralogy, 10, 395–406.
Kirby, S.H., 1983. Rheology of the lithosphere. Reviews of Geophysics, 21, 1458–1487.
Li, Z.H., Gerya, T.V., and Burg, J.P., 2010. Influence of tectonic overpressure on P–T paths of HP–
UHP rocks in continental collision zones: thermomechanical modelling. Journal of
Metamorphic Geology, 28, 227–247.
16
Liou, J.G., Hacker, B.R., and Zhang, R.Y., 2000. Into the Forbidden Zone. Science 287, 1215–
1216.
McKenzie, D., Jackson, J., and Priestley, K., 2005. Thermal structure of oceanic and continental
lithosphere. Earth and Planetary Science Letters, 233, 337–349.
Ranalli, G., 1995. Rheology of the Earth. Chapman & Hall, London.
Ranalli, G., and Murphy, D.C., 1987. Rheological stratification of the lithosphere. Tectonophysics,
132, 281–295.
Schilling, J.-G., Zajec, M., Evans, R., Johnston, T., White, W., Devine, J. D., and Kingsley, R.,
1983. Petrologic and geochemical variations along the Mid-Atlantic Ridge from 27ºN to
73ºN. American Journal of Science 283, 510–586.
Shellnutt, J.G., Bhat, G.M., Wang, K-L., Brookfield, M.E., Jahn, B.M., and Dostal, J., 2014.
Petrogenesis of the flood basalts from the Early Permian Panjal Traps, Kashmir, India:
Geochemical evidence for shallow melting of the mantle. Lithos, 204, 159–171.
Shelton, G., and Tullis, J., 1981. Experimental flow laws for crustal rocks. Eos Transactions,
American Geophysical Union, 62, 396.
Syracuse, E.M., van Keken, P.E., and Abers, G.A., 2010. The global range of subduction zone
thermal models. Physics of the Earth and Planetary Interiors, 183, 73–90.
Turcotte, D.L., and Schubert, G., 2014. Geodynamics: Cambridge University Press.
White, R.W., Powell, R., Holland, T.J.B., and Worley, B.A., 2000. The effect of TiO2 and Fe2O3 on
metapelitic assemblages at greenschist and amphibolite facies conditions: mineral equilibria
calculations in the system K2O–FeO–MgO–Al2O3–SiO2–H2O–TiO2–Fe2O3. Journal of
Metamorphic Geology, 18, 497–511.
White, R.W., Powell, R., and Holland, T.J.B., 2007. Progress relating to calculation of partial
melting equilibria for metapelites. Journal of Metamorphic Geology, 25, 511–527.
Top Related