Crustal evolution and mantle dynamics through Earth history

Abstract

Resolving the modes of mantle convection through Earth history, i.e. when plate tectonics started and what kind of mantle dynamics reigned before, is essential to the understanding of the evolution of the whole Earth system, because plate tectonics influences almost all aspects of modern geological processes. This is a challenging problem because plate tectonics continuously rejuvenates Earth's surface on a time scale of about 100 Myr, destroying evidence for its past operation. It thus becomes essential to exploit indirect evidence preserved in the buoyant continental crust, part of which has survived over billions of years. This contribution starts with an in-depth review of existing models for continental growth. Growth models proposed so far can be categorized into three types: crust-based, mantle-based and other less direct inferences, and the first two types are particularly important as their difference reflects the extent of crustal recycling, which can be related to subduction. Then, a theoretical basis for a change in the mode of mantle convection in the Precambrian is reviewed, along with a critical appraisal of some popular notions for early Earth dynamics. By combining available geological and geochemical observations with geodynamical considerations, a tentative hypothesis is presented for the evolution of mantle dynamics and its relation to surface environment; the early onset of plate tectonics and gradual mantle hydration are responsible not only for the formation of continental crust but also for its preservation as well as its emergence above sea level. Our current understanding of various material properties and elementary processes is still too premature to build a testable, quantitative model for this hypothesis, but such modelling efforts could potentially transform the nature of the data-starved early Earth research by quantifying the extent of preservation bias.This article is part of a discussion meeting issue 'Earth dynamics and the development of plate tectonics'.

Keywords: continental growth; mantle convection; plate tectonics.

Figure 1.

Various published models for the growth of continental crust: HR67 [13], F78 [22], B79 [23], A81 [24], MT82 [25], RS84 [26], AR84 [16], ND85 [27], PA86 [28], J88 [29], W89 [30], CK99 [31], C03 [32], R04 [33], CA10 [34], B10 [35], D12 [11], K18 [36] and RK18 [37]. The volume or mass of the continental crust is normalized by the present-day value. Line colours are assigned to produce gradations from blue to red across the models; each model retains its line colour in figures 2, 3 and 6.

Figure 2.

Crust-based ‘growth’ models discussed in §2a. As indicated here, these models aim to estimate on the present-day distribution of formation ages, so they serve as the lower bound on net crustal growth. The cumulative distribution is normalized by the present-day continental mass or area. Also shown in dashed line is the surface age distribution according to Goodwin [41].

Figure 3.

Mantle-based models of net crustal growth discussed in §2b.

Figure 4.

(a) Schematic diagram to illustrate the relation among net crustal growth (from mantle-based models), the formation age distribution (from crust-based models), the surface age distribution, crustal recycling and crustal reworking. In particular, the difference between crust- and mantle-based models reflects the extent of crustal recycling, thus the operation of plate tectonics. (b) Before 2018, the difference between the latest crust- and mantle-based models was remarkably small. Note that the model of Dhuime et al. [11] cannot be a model for net crustal growth because of the nature of their data. (c) The latest crust-based model [36] and the latest mantle-based model [37] as of 2018. Rosas & Korenaga [37] employed Monte Carlo sampling, and dark and light shades denote the 50% and 90% confidence limits, respectively, of the successful Monte Carlo ensemble. The Armstrong model [24] is also shown for comparison.

Figure 5.

(a) Crustal recycling rate (orange shades) corresponding to the net crustal growth model of Rosas & Korenaga [37]. The meaning of dark and light shades is the same as in figure 4c. For comparison, the present-day recycling rate based on sediment subduction is 0.7–0.9 × 10²² kg Gyr⁻¹ [85,86]. The rate of new crustal generation is shown in purple dashed line. For the sake of clarity, only the median solution is shown for the crustal generation rate. (b) The present-day formation age distribution predicted from the box modelling of the Nd isotope evolution [37] (blue shades) is compared with that estimated from the global database of detrital zircon [36] (dark blue line).

Figure 6.

Crustal growth models in the third category (i.e. neither crust-based nor mantle-based) discussed in §2d.

Figure 7.

Hypothetical yield strength profiles for present-day oceanic lithosphere at the seafloor age of (a) 30 Ma and (b) 100 Ma. Only the top 100 km is shown, and the assumed thermal structure, based on half-space cooling with the initial temperature of 1350°C, is indicated on the right axis. Yield strength is calculated assuming the geological strain rate of 10⁻¹⁵ s⁻¹, so the yield strength of 1 GPa corresponding to the effective viscosity of 10²⁴ Pa s. The effect of thin crustal layer is ignored, and deformation mechanisms considered here include: (1) diffusion creep with the activation energy of 300 kJ mol⁻¹ and the grain size exponent of 2, with the reference viscosity of 10¹⁹ Pa s at 1350°C and the grain size of 5 mm, (2) dislocation creep with the activation energy of 600 kJ mol⁻¹ and the stress exponent of 3, with the reference viscosity of 10¹⁹ Pa s at 1350°C and the deviatoric stress of 0.1 MPa, (3) low-temperature plasticity based on the reanalysis of the experimental data of Mei et al. [123] by Jain et al. [124], with the exponents of p = 1 and q = 2, (4) linear exponential viscosity with the total viscosity contrast of 10⁵, (5) brittle strength with the friction coefficient of 0.8 and with optimal thrust faulting [125] and (6) brittle strength with the effective friction coefficient of 0.03 to emulate the effect of thermal cracking [126]. Shown for comparison is the constant yield strength of 150 MPa, which is a typical value used in numerical simulation studies. Diffusion creep with the grain size of 10 mm is shown to illustrate the efficacy of grain-size reduction; this weakening mechanism is not effective in the cold part of lithosphere. Thick dashed lines trace two contrasting cases: case A (pink, short-dashed) for the combination of linear exponential viscosity with constant yield strength, and case B (blue, long-dashed) for the combination of more realistic ductile deformation mechanisms and thermal cracking. These cases are different by up to two orders of magnitudes over the bulk of lithosphere.

Figure 8.

(a) The thickness of top thermal boundary layer as a function of mantle potential temperature [117]. For the calculation of convective instability, the reference viscosity of 10¹⁹ Pa s at 1350°C and the activation energy of 300 kJ mol⁻¹ are used. The case of purely temperature-dependent viscosity is shown in red, and the case with the effect of mantle melting (on viscosity as well as density) in blue. Assumed dry and wet solidi are also shown. (b) Surface heat flow as a function of mantle potential temperature [117], based on the scaling of boundary layer thickness shown in (a), and assuming the operation of plate tectonics. Star denotes the present-day condition. Both scaling laws were originally derived from a simple boundary layer theory, but their first-order features are also supported by fully dynamic convection models [126,138].

Figure 9.

(a) Petrological estimates on the past potential temperature of the ambient mantle (grey circles) [97], and the predicted thermal history (blue curve) using the constant surface heat flow with the present-day Urey ratio of 0.22 [134]. The core heat flow is assumed to vary linearly from 15 TW at 4.5 Ga to 10 TW at present [172]. (b) Surface heat flow (blue), core heat flow (yellow) and radiogenic heat production in the mantle (red) used for the thermal history shown in (a). Also shown is corresponding plate velocity (green), with the present-day average velocity of 5 cm yr⁻¹ [173].

Figure 10.

Schematic for the working hypothesis on the coupled ocean–crust–mantle evolution. The illustrations are not drawn to scale in every possible detail, though certain details are intentional. The topography of continental crust was probably more reduced in the past because of hotter geotherm [213]. Continental mantle lithosphere could have been thicker in the past if we take into account the effect of convective erosion [208,214]. No significant difference in plate velocity is expected (figure 9b), but the structure of oceanic lithosphere is very different because of the secular cooling of the mantle (figures 8a and 9a). The average thickness of continental crust is likely to have been stable since around 3 Ga [215]. Continental crust could have been largely submarine during the Archean, as suggested by the abundance of submarine continental flood basalts [216] and also implied by freeboard modelling [100].