Nature | Vol 586 | 15 October 2020 | 397not descend and detaches easily from the sinking, cooler and more rigid
subducting slab. The ascending, depleted and low-density mantle origi-
nally accumulates as a bulge under the bend in the subducting oceanic
plate (Fig. 1a). The growing volume of this bulge triggers relatively rapid
(similar to the rate of convergence) lateral viscous flow directed away
from the trench towards the incoming continental margin, where it is
emplaced laterally as a thick, depleted sub-lithospheric layer (proto-
keel) under the continent, thus creating chemical and rheological layer-
ing in the subcontinental mantle (Fig. 1c, Extended Data Fig. 2). Density
profiles across the protokeel section reveal a layer about 50 km thick
of low-density mantle at depths of 145–190 km (Fig. 1d, inset). Owing
to the high ambient mantle temperature and large slab-pull forces
produced by a thick and dense eclogitized oceanic crust, the subduct-
ing slabs experience frequent episodes of slab break-off^25 , thereby
returning detached fragments of the slab to the mantle transition zone
(Fig. 1d). During the subsequent arc-continent collision, the supply of
hot (>1,300 °C) depleted material to the viscous underplate vanishes,
and the protokeel volume stops growing (Fig. 1d, Extended Data Fig. 3).
At this stage, the underplate may also almost lose connection with the
cold (<1,300 °C) downgoing rigid mantle lithosphere (Fig. 1d, Extended
Data Fig. 3b). This is caused by the disappearance of the source of
depleted oceanic sub-lithospheric mantle after complete consump-
tion of the oceanic plate, since such hot depleted mantle is not present
under the cold subducting continental lithosphere.
The thickness of the ductile, low-density, melt-depleted sub-
lithospheric mantle present under the oceanic plate depends mainly
on the mantle potential temperature. The higher the temperature,
the larger the volume of melt-depleted ductile mantle arriving at the
subduction zone and contributing to the formation of the viscous
protokeel. In our models, the largest/longest/deepest viscous proto-
keels develop at the highest modelled mantle potential temperature
(ΔT = 250 °C, Tp = 1,550 °C, Fig. 2a, b), whereas they almost disappear
when this temperature is still 150 °C higher than the present-day values
(ΔT = 150 °C, Tp = 1,450 °C, Fig. 2c, d). The size of the viscous proto-
keels should probably correlate with the volume of oceanic lithosphere
involved in the subduction. The characteristic length of lithospheric
blocks detached by slab break-off also depends on the mantle tempera-
ture: the higher the temperature, the more frequent are the break-offs
and the smaller are the blocks stored in the mantle transition zone
(Fig. 1d and Fig. 2b, d). Owing to very rapid subduction, some amount of
water remains preserved in the eclogitized crust of these blocks. Nota-
bly, such blocks also contribute to the formation of sub-continental
mantle as numerous small mantle diapirs originate from hydrated,
depleted slab fragments present in the mantle transition zone (Fig. 1d
and Fig. 2b, d). This numerical-modelling prediction agrees well with
previous studies that investigated deep stagnant slab dehydration and
the formation of hydrous diapirs in the mantle transition zone^26 ,^27. The
diapirs detach from the dense (eclogitized) subducted oceanic crust
and their positive buoyancy is related to both the initial depletion of the
subducted lithospheric mantle and to the hydration by fluids derived
from the crust.Viscous underplates in cratonic keels
The structure of lithospheric keels can be evaluated by reference to
‘chemical tomography’ sections^28 , derived using garnet and chromite
xenocrysts from kimberlites and related rocks worldwide^14 (see Meth-
ods). Many of these sections, such as the Daldyn Field of the Siberian cra-
ton in eastern Russia; the Slave craton in northern Canada; the Limpopo
Belt, north Lesotho and north Botswana in southern Africa; and the
Gawler Craton in southern Australia, show clear compositional strati-
fication (Extended Data Figs. 4–7) generalized in Fig. 3a, b. TypicallyΔT = 250 ºC
8.2 MyrΔT = 150 ºC
6.3 MyrΔT = 150 ºC
9.6 MyrΔT = 250 ºC
16.4 MyrViscous ow of
depleted mantle5 ɫm yr–1Continental crust Oceanic crust Continental crust Oceanic crustStagnant slabs Mantle transition
zone
Stagnant
slab5 ɫm yr–1DiapirsDiapirsabcd1,800 2,000 2,200 2,400 2,600 2,800 3,000 1,800 2,000 2,200 2,400 2,600 2,800 3,00020060040005003001002004000300100Mantle MantleDistance (km)ViDepleted mantle Depleted mantleMantle transition
zoneDepth (km)Depth (km)Distance (km)protokeelMantleFig. 2 | Effect of mantle temperature on the development of mantle keels in
subduction/collision zones. a, b, Development of 1,000-km-long viscous-mantle
underplate and a number of small stagnant slabs in the mantle transition zone
at elevated mantle temperature ΔT = 250 °C (Tp = 1, 550 °C) at 8.2 Myr (a) and
16.4 Myr (b) from the beginning of the experiment. c, d, Evolution of subduction/
collision zone with incipient viscous underplate and large stagnant slab in a
transition zone at elevated mantle temperature ΔT = 150 °C (Tp = 1,450 °C) at
6.3 Myr (c) and 9.6 Myr (d) from the beginning of the experiment. Dotted yellow
lines indicate upper and lower boundaries of the mantle transition zone.
Arrows show direction of plate motion and viscous f low of the depleted mantle.
The colour key is as in Fig. 1.