The Tibetan Plateau is the largest (~1,500 × 3,500 km), high-elevation (mean of ~5,000 m) topographic
feature on Earth and hosts the thickest crust of any modern orogen, with estimates in southern Tibet of
~70 km (Owens and Zandt, 1997; Nábělek et al., 2009), and up to ~85 km (Wittlinger et al., 2004; Xu et
al., 2015). The Tibetan Plateau formed from the sequential accretion of continental fragments and island
arc terranes beginning during the Paleozoic and culminated with the Cenozoic collision between India and
Asia (Argand, 1922; Yin and Harrison, 2000; Kapp and DeCelles, 2019). The India-Asia collision is
largely thought to have commenced between 60 and 50 Ma (e.g., Rowley, 1996; Hu et al., 2016); however,
some raise the possibility for later collisional onset (e.g., Aitchison et al., 2007; van Hinsbergen et
al., 2012). Despite ongoing ~north-south convergence, the northern Himalaya and Tibetan Plateau interior
are undergoing east-west extension, expressed as an array of approximately north-trending rifts that
extend from the axis of the high Himalayas to the Bangong Suture Zone (Molnar and Tapponnier, 1978;
Taylor and Yin, 2009) (Fig. 1).
Digital elevation model of southern Tibet with major tectonic features. Active structures from
(Styron et al., 2010). The basemap is from MapBox Terrain Hillshade. Lake locations are
from Yan et al. (2019). Data points include only the filtered data (supplemental material [see text
footnote 1]). HW—hanging wall.
The Mesozoic tectonic evolution of the southern Asian margin placed critical initial conditions for the
Cenozoic evolution of the Tibetan Plateau. However, much of the Mesozoic geologic history remains poorly
understood, in part due to structural, magmatic, and erosional modification during the Cenozoic. There
is disagreement even on first-order aspects of the Mesozoic geology in the region. For example, temporal
changes in Mesozoic crustal thickness are largely unknown, and the paleoelevation of the region is
debated. Most tectonic models invoke major shortening and crustal thickening due to shallow subduction
during the Late Cretaceous (e.g., Wen et al., 2008; Guo et al., 2013), possibly pre-conditioning the
southern Asian margin as an Andean-style proto-plateau (Kapp et al., 2007; Lai et al., 2019).
Alternatively, Late Cretaceous to Paleogene shortening may have been punctuated by a 90–70 Ma phase of
extension that led to the rifting of a southern portion of the Gangdese arc and opening of a backarc
ocean basin (Kapp and DeCelles, 2019). These represent two competing end-member hypotheses for the
Mesozoic tectonic evolution of southern Tibet that are testable by answering the question: Was the crust
in southern Tibet thickening or thinning between 90 and 70 Ma?
Contrasting hypotheses about the Cenozoic tectonic evolution of southern Tibet are testable by
quantifying changes in crustal thickness through time. In particular, the Paleocene tectonic evolution
before, during, and after the collision between India and Asia was dependent on initial crustal
thickness, and in part controlled the development of the modern Himalayan-Tibetan Plateau. Building on
the hypothesis tests for the Late Cretaceous, if the crust of the southern Asian margin was thickened
before or during the Paleocene, then this explains why the southern Lhasa Terrane was able to attain
high elevations only a few million years after the onset of continental collisional orogenesis (Ding et
al., 2014; Ingalls et al., 2018). However, if the Paleocene crust was thin, then we can ask the
question: When did the crust attain modern or near modern thickness? Answering this question is a
critical test of alternative tectonic models that suggest rapid surface uplift from relatively low
elevation (and presumably thin crust) during the Miocene (e.g., Harrison et al., 1992; Molnar et al.,
1993) or Pliocene (Dewey et al., 1988) as the product of mantle lithosphere removal (England and
Houseman, 1988). Finally, what happened after the crust was thickened to extreme levels, as we have in
the modern? Did the plateau begin to undergo orogenic collapse (Dewey, 1988) resulting in a net
reduction in crustal thickness and surface elevation that continues to present day (e.g., Ge et al.,
2015), as evidenced by the Miocene onset of east-west extension (e.g., Harrison et al., 1995; Kapp et
al., 2005; Sanchez et al., 2013; Styron et al., 2013, 2015; Sundell et al., 2013; Wolff et al., 2019)?
Or did Tibet remain at steady-state elevation during Miocene-to-modern extension (Currie et al., 2005)
with upper crustal thinning and ductile lower crustal flow (e.g., Royden et al., 1997) working to
balance continued crustal thickening at depth driven by the northward underthrusting of India (DeCelles
et al., 2002; Kapp and Guynn, 2004; Styron et al., 2015)?
Igneous rock geochemistry has long been used to estimate qualitative changes in past crustal (e.g.,
Heaman et al., 1990) and lithospheric (e.g., Ellam, 1992) thickness. Trace-element abundances of igneous
rocks have proven particularly useful for tracking changes in crustal thickness (Kay and Mpodozis, 2002;
Paterson and Ducea, 2015). Trace-element ratios provide information on the presence or absence of
minerals such as garnet, plagioclase, and amphibole because their formation is pressure dependent, and
each has an affinity for specific trace elements (e.g., Hildreth and Moorbath, 1988). For example, Y and
Yb are preferentially incorporated into amphibole and garnet in magmatic melt residues, whereas Sr and
La have a higher affinity for plagioclase (Fig. 2A). Thus, high Sr/Y and La/Yb can be used to infer a
higher abundance of garnet and amphibole and a lower abundance of plagioclase, and may be used as a
proxy for assessing the depth of parent melt bodies during crustal differentiation in the lower crust
(Heaman et al., 1990). These ratios have been calibrated to modern crustal thickness and paired with
geochronological data to provide quantitative estimates of crustal thickness and paleoelevation through
time (e.g., Chapman et al., 2015; Profeta et al., 2015; Hu et al., 2017, 2020; Farner and Lee, 2017).
(A) Schematic partitioning diagram for Y and Yb into minerals stable at high lithostatic pressures >1
GPa such as garnet and amphibole. (B–D) Empirical calibrations using known crustal thicknesses from data
compiled in Profeta et al. (2015) based on (B) multiple linear regression of ln(Sr/Y) (x
-axis), and crustal thickness (z
-axis); (C) simple linear regression of
ln(Sr/Y) and crustal thickness; and (D) simple linear regression of ln(La/Yb) and crustal thickness.
Equations in parts B–D include 95% confidence intervals for each coefficient. Coefficient uncertainties
should not be propagated when applying these equations to calculate crustal thickness; rather, the 2s
(95% confidence interval) residuals (modeled fits subtracted from known crustal thicknesses) are more
representative of the calibration uncertainty.
We build on recent efforts to empirically calibrate trace-element ratios of igneous rocks to crustal
thickness and apply these revised calibrations to the eastern Gangdese mountains in southern Tibet (Fig.
1). This region has been the focus of several studies attempting to reconstruct the crustal thickness
using trace-element proxies (e.g., Zhu et al., 2017) as well as radiogenic isotopic systems such as Nd
and Hf (Zhu et al., 2017; Alexander et al., 2019; DePaolo et al., 2019), and highlight discrepancies in
different geochemical proxies of crustal thickness. As such, we first focus on developing a new approach
to estimate crustal thickness from Sr/Y and La/Yb, both for individual ratios, and in paired Sr/Y–La/Yb
calibration. We then apply these recalibrated proxies to data from the Gangdese mountains to test
hypotheses explaining the Mesozoic and Cenozoic tectonic evolution of southern Tibet.
Sr/Y and La/Yb (the latter normalized to the chondritic reservoir) were empirically calibrated using a
modified approach reported in Profeta et al. (2015). Calibrations are based on simple linear regression
of ln(Sr/Y)–km and ln(La/Yb)–km; and multiple linear regression of ln(Sr/Y)–ln(La/Yb)–km (Figs. 2B–2D).
We also tested simple linear regression of ln(Sr/Y) × ln(La/Yb)–km (see GSA Supplemental
Material1). Regression coefficients and residuals (known minus modeled thickness) are
reported at 95% confidence (±2s).
The revised proxies were applied to geochemical data compiled in the Tibetan Magmatism Database (Chapman
and Kapp, 2017). Geochemical data used here comes from rocks collected in an area between 29 and 31°N
and 89 and 92°E. Data were filtered following methods reported in Profeta et al. (2015) where samples
outside compositions of 55%–68% SiO2, 0%–4% MgO, and 0.05–0.2% Rb/Sr are excluded to avoid
mantle-generated mafic rocks, high-silica felsic rocks, and rocks formed from melting of metasedimentary
rocks. Filtering reduced the number of samples considered from 815 to 190 (supplemental material; see
We calculated temporal changes in crustal thickness based on multiple linear regression of
ln(Sr/Y)–ln(La/Yb)–km (Fig. 2B). Each estimate of crustal thickness is assigned uncertainty of ±5 m.y.
and ±10 km; the former is set arbitrarily because many samples in the database do not have reported
uncertainty, and the latter is based on residuals calculated during proxy calibration (Fig. 2). Temporal
trends were calculated using two different methods. The first method employs Gaussian kernel regression
(Horová et al., 2012), a non-parametric technique commonly used to find nonlinear trends in noisy
bivariate data; we used a 5 m.y. kernel width, an arbitrary parameter selected based on sensitivity
testing for over- and under-smoothing. The second method involves calculating linear rates between
temporal segments bracketed by clusters of data that show significant changes in crustal thickness:
200–150 Ma, 100–65 Ma, and 65–30 Ma. Trends are reported as the mean ±2s calculated from bootstrap
resampling 190 selections from the data with replacement 10,000 times.
Proxy calibration using simple linear regression of ln(Sr/Y)–km and ln(La/Yb)–km yields
Crustal Thickness = (19.6 ± 4.3) × ln(Sr/Y) + (−24.0 ±
Crustal Thickness = (17.0 ± 3.7) × ln(La/Yb) + (6.9 ±
whereas multiple linear regression of ln(Sr/Y)–ln(La/Yb)–km calibration yields
Crustal Thickness = (−10.6 ± 16.9) + (10.3 ± 9.5) × ln(Sr/Y) + (8.8 ± 8.2) × ln(La/Yb). (3)
Crustal thickness corresponds to the depth of the Moho in km, and coefficients are ±2s (Figs. 2B–2D and
supplemental material [see footnote 1]). Although we report uncertainties for the individual
coefficients, propagating these uncertainties results in wildly variable (and often unrealistic) crustal
thickness estimates, largely due to the highly variable slope. Hence, we ascribe uncertainties based on
the 2s range of residuals (Figs. 2B–2D). Residuals are ~11 km based on simple linear regression of
Sr/Y–km and La/Yb–km, and ~8 km based on multiple linear regression of Sr/Y–La/Yb–km.
Application of these equations yields mean absolute differences between crustal thicknesses calculated
with individual Sr/Y and La/Yb of ~6 km. Paired Sr/Y–La/Yb calibration yields absolute differences of ~3
km compared to Sr/Y and La/Yb. Discrepancies in crustal thickness estimates between Sr/Y and La/Yb using
the original calibrations in Profeta et al. (2015) are highly variable, with an average of ~21 km, and
are largely the result of extreme crustal thickness estimates (>100 km) resulting from linear
transformation of high (>70) Sr/Y ratios (supplemental material [see footnote 1]); such discrepancies
are likely due to a lack of crustal thickness estimates from orogens with rocks that are young enough
(i.e., Pleistocene or younger) to include in the empirical calibration.
For geologic interpretation, we use results from multiple linear regression of Sr/Y–La/Yb–km to calculate
temporal changes in crustal thickness (Figs. 3B–3D). Results show a decrease in crustal thickness from
36 to 30 km between 180 and 170 Ma. Available data between 170 and 100 Ma include a single estimate of
~55 km at ca. 135 Ma. Crustal thickness decreased to 30–50 km by ca. 60 Ma, then increased to 60–70 km
by ca. 40 Ma (Fig. 3). The two different methods for calculating temporal trends in crustal thickness
(Gaussian kernel regression and linear regression) produced similar results (Figs. 3C–3D). The Gaussian
kernel regression model produces a smooth record of crustal thickness change that decreases from ~35 to
~30 km between 180 and 165 Ma, decreases from ~54 to ~40 km between 90 and 75 Ma, increases from ~40 to
~70 km between 60 and 40 Ma, and remains steady-state from 40 Ma to present; the large uncertainty
window between 160 and 130 Ma is due to the bootstrap resampling occasionally missing the single data
point at ca. 135 Ma (Fig. 3C). Linear rates of crustal thickness change indicate thinning at ~0.7 mm/a
between 180 and 170 Ma, thinning at ~0.8 mm/a between 90 and 65 Ma, and thickening at ~1.3 mm/a between
60 and 30 Ma (Fig. 3D).
Results of new Sr/Y and La/Yb proxy calibration applied to data from the Tibetan Magmatism Database
(Chapman and Kapp, 2017) located in the eastern Gangdese Mountains in southern Tibet. (A) Filtered Sr/Y
and La/Yb data extracted from 29 to 31°N and 89 to 92°W. (B) Values from part A converted to crustal
thickness using Equations 1, 2, and 3 (see text). (C–D) Temporal trends based on multiple linear
regression of Sr/Y–La/Yb–km; trends are calculated from 10,000 bootstrap resamples, with replacement.
(C) Gaussian kernel regression model to determine a continuous thickening history. (D) Linear regression
to determine linear rates for critical time intervals.
Early to Middle Jurassic crustal thickness in southern Tibet was controlled by the northward subduction
of Neo-Tethys oceanic lithosphere (Guo et al., 2013) in an Andean-type orogen that existed until the
Early Cretaceous (Zhang et al., 2012), punctuated by backarc extension between 183 and 174 Ma (Wei et
al., 2017). The latter is consistent with our results of minor crustal thinning from ~36 to ~30 km
between 180 and 170 Ma (Figs. 3B–3D) and supports models invoking a period of Neo-Tethys slab rollback
(i.e., trench retreat), southward rifting of the Zedong arc from the Gangdese arc, and a phase of
supra-subduction zone ophiolite generation along the southern margin of Asia (Fig. 4A) (Kapp and
DeCelles, 2019). Rocks with ages between 170 and 100 Ma are limited to a single data point at ca. 135 Ma
and yield estimates of ~55 km thick crust (Fig. 3B–3D). This is consistent with geologic mapping and
geochronological data that suggest that major north-south crustal shortening took place in the Early
Cretaceous along east-west–striking thrust faults in southern Tibet (Murphy et al., 1997).
Tectonic interpretation after Kapp and DeCelles (2019). (A) Middle–Late Jurassic accelerated slab
rollback during formation of the Zedong Arc drives the opening of an extensional backarc basin. This is
consistent with the generation of the late-stage, juvenile (asthenosphere derived) Yeba volcanics (Liu
et al., 2018). (B) Late Cretaceous slab rollback results in the opening of a backarc ocean basin.
The strongest crustal thinning trend in our results occurs between 90 and 70 Ma at a rate of ~0.8 mm/a
(Fig. 3D). Crustal thinning takes place after major crustal shortening (and thickening) documented in
the southern Lhasa terrane prior to and up until ca. 90 Ma (Kapp et al., 2007; Volkmer et al., 2007; Lai
et al., 2019), following shallow marine carbonate deposition during the Aptian–Albian (126–100 Ma)
across much of the Lhasa terrane (Leeder et al., 1988; Leier et al., 2007). Late Cretaceous crustal
thinning to ~40 km (closer to the average thickness of continental crust) supports models that invoke
Late Cretaceous extension and Neo-Tethys slab rollback that led to the development of an
intracontinental backarc basin in southern Tibet and southward rifting of a southern portion of the
Gangdese arc (referred to as the Xigaze arc) from the southern Asian continental margin (Kapp and
DeCelles, 2019) (Fig. 4B). If a backarc ocean basin indeed opened between Asia and the rifted Xigaze arc
during this time, it would have profound implications for Neo-Tethyan paleogeographic reconstructions
and the history of suturing between India and Asia; this remains to be tested by future field-based
Paleogene crustal thickness estimates indicate monotonic crustal thickening at rates of ~1.3 mm/a to
>60 km following the collision between India and Asia. This is in contrast to models explaining the
development of modern high elevation resulting from the removal of mantle lithosphere during the Miocene
or Pliocene (Dewey et al., 1988; England and Houseman, 1988; Harrison et al., 1992; Molnar et al.,
1993). The timing of crustal thickening in the late Paleogene temporally corresponds to the termination
of arc magmatism in southern Tibet at 40–38 Ma and may indicate that the melt-fertile upper-mantle wedge
was displaced to the north by shallowing subduction of Indian continental lithosphere (Laskowski et al.,
2017). Crustal thickening during the Paleogene may be attributed to progressive shortening and southward
propagation (with respect to India) of the Tibetan-Himalayan orogenic wedge as Indian crust was accreted
in response to continuing convergence. We interpret that thickening depended mainly on the flux of crust
into the orogenic wedge, as convergence between India and Asia slowed by more than 40% between 20 and 10
Ma (Molnar and Stock, 2009), subsequent to peak crustal thickening rates between 60 and 30 Ma.
Estimates of crustal thickness based on Sr/Y and La/Yb differ both in time and space compared to
estimates using radiogenic isotopes. Determining crustal thickness from Nd or Hf relies on an extension
of the flux-temperature model of DePaolo et al. (1992), which calculates the ambient crustal temperature
and assimilation required to produce measured isotopic compositions (e.g., Hammersley and DePaolo, 2006)
assuming a depleted asthenospheric melt source with no contribution from the mantle lithosphere; crustal
thickness is then calculated based on an assumed geothermal gradient on the premise that a deeper,
hotter Moho would result in more crustal assimilation than a shallower, cooler Moho. In addition to
using La/Yb to estimate Cenozoic crustal thickness, DePaolo et al. (2019) use the flux-temperature model
to suggest that crustal thickening in southern Tibet was nonuniform based on Nd isotopes. Specifically,
they estimate crustal thickness of 25–35 km south of 29.8° N until 45 Ma, followed by major crustal
thickening to 55–60 km by the early to middle Miocene. Critically, they suggest that north of 29.9° N
the crust was at near modern thickness before 45 Ma and that there was a crustal discontinuity between
these two domains, which Alexander et al. (2019) later interpret along orogenic strike to the east based
on Hf isotopic data. In contrast, our results show that crustal thickening was already well under way by
45 Ma, potentially near modern crustal thickness, and with no dependence on latitude (Figs. 3B–3D and
supplemental material [see footnote 1]). Radiogenic isotopes such as Nd and Hf are not directly
controlled by crustal thickness and concomitant pressure changes. Rather, variability in Hf and Nd is
likely due to complex crustal assimilation, to which pressure-based (not temperature-based) proxies such
as Sr/Y and La/Yb from rocks filtered following Profeta et al. (2015) are less sensitive.
Crustal thickness of 65–70 km between 44 and 10 Ma based on trace-element geochemistry is similar to
modern crustal thickness of ~70 km estimated from geophysical methods (Owens and Zandt, 1997; Nábělek et
al., 2009) and are 10–20 km less than upper estimates of 80–85 km (Wittlinger et al., 2004; Xu et al.,
2015). Upper-crustal shortening persisted in southern Tibet until mid-Miocene time, but coeval rapid
erosion (Copeland et al., 1995) may have maintained a uniform crustal thickness. Our results are
inconsistent with models that invoke net crustal thinning via orogenic collapse (Dewey, 1988) beginning
in the Miocene and continuing to present day (Ge et al., 2015). Rather, our results are consistent with
interpretations of thick crust in southern Tibet by middle Eocene time (Aikman et al., 2008; Pullen et
al., 2011), which continued to thicken at depth due to the ongoing mass addition of underthrusting India
(DeCelles et al., 2002) before, during, and after the Miocene onset of extension in southern Tibet
(e.g., Harrison et al., 1995; Kapp et al., 2005; Sanchez et al., 2013). We favor a model in which
continued crustal thickening at depth is balanced by upper crustal thinning (Kapp and Guynn, 2004;
DeCelles et al., 2007; Styron et al., 2015), with excess mass potentially evacuated by ductile lower
crustal flow (Royden et al. 1997). In this view, late Miocene–Pliocene acceleration of rifting in
southern Tibet (Styron et al., 2013; Sundell et al., 2013; Wolff et al., 2019) is a consequence of the
position of the leading northern tip of India (Styron et al., 2015), because this region experiences
localized thickening at depth, which in turn increases the rate of upper crustal extension in order to
maintain isostatic equilibrium.
We thank Chris Hawkesworth, Allen Glazner, two anonymous reviewers, and editor Peter Copeland for their
detailed critique of this work. We also thank Michael Taylor and Richard Styron for informal reviews;
Sarah George and Gilby Jepson for insightful discussions on proxies for crustal thickness; and Caden
Howlett and Aislin Reynolds for discussions of Tibetan tectonics during the 2019 field season as this
research was formulated. KES was partially supported by the National Science Foundation (EAR-1649254) at
the Arizona LaserChron Center. M.N.D. acknowledges support from National Science Foundation grant EAR
1725002 and the Romanian Executive Agency for Higher Education, Re-search, Development and Innovation
Funding project PN-III-P4-ID-PCCF-2016-0014.
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