9. Questions
• How has magmatism evolved through Earth
history?
• How has continental tectonics evolved
through Earth history?
• How have these evolutions controlled the
chemistry of the ocean and atmosphere ?
12. Xe isotopes tell us how long the magma
« era » has lasted?
(Kunz, Staudacher, Allègre 1998)
8.0
7.8
7.6
7.4
129
I
Xe/130Xe
7.2
7.0
129
6.8
atmosphere
6.6
244
Pu + 238U
6.4
6.2 atm. non rad.
6.0 Half-lives :
2.00 2.10 2.20 2.30
136
2.40 2.50
129
I
2.60
16 Ma
2.70
Xe/130Xe
244
Pu 82 Ma
238
U 4450 Ma
13. Undegassed
mantle Present-day
6800
100 mantle
2-22
10
30
0.25-1.5
1 1
1 1 2 3
Xe
136
Xe
136
Xe
129
Xe
136
Xe
136
Xe
129
238
U 244
Pu 129
I 238
U 244
Pu 129
I
14. Mantle degassing
• Accretion (impacts)
Complete and instantaneous
• Magma ocean (melting)
• Secular cooling (convection)
Incomplete and slow (~ 4.4By)
27. Subsidence
3 Myrs after loading
°C °C
= 5 00 = 8 00
T mo
ho T mo ho
Flament et al., 2011
28. Submarine volcanism during the Archean
Archean (~3.5 Ga) pillow lava (P. Rey) Kump and Barley (2007)
Flood volcanism on submerged continental platforms is:
• common in the Precambrian
• rare to absent throughout the Phanerozoic
These lavas have interacted with differentiated material
(N. Arndt, 1999)
34. Proposed synthesis
• Hadean: « magma » tectonics and impacts combined to
high temperature hydrothermalism : buffered soda
ocean.
• Archean: plate tectonics (?) with soft continents
combined to hydrothermalism: from a soda to halite
ocean?
• Post-Archean: plate tectonics with stiff continents
combined to weathering and hydrothermalism: modern
ocean.
Editor's Notes
The evolution of the atmosphere and oceans is intimately linked to the evolution of the deep interior of the Earth. The coupling between the thin fluid enveloppe and the thick rocky mantle occurs through different processes that have changed over time, and especially in the first 2 billion years of the planetary evolution. The size of the fluid and silicate reservoirs are fundamentally different. Hence, the surface/volume ratio is maximum for the oceans, but very limited from the point of view of the mantle and crust.
Carbon is a focus of study for the interaction between the deep and surficial Earth, because it can be solid, gaseous or dissolved and it has a huge impact on the environment and life development. Here the long-term carbon cycle is described for a modern Earth. Transfers of carbon between solid and fluid enveloppes clearly regulate the whole cycle. In their paper, Sleep and Zahnle suggest that this cycle has not remained identical through the Earth history. Three main processes of mass exchange between the exosphere, the continents and the mantle operate.
The first process is degassing in volcanic areas: decompression of the magma generates the exolution of gases and huge amounts of gases and especially CO2 are released. Degassing is responsible for most of the atmospheric inventory of volatiles. CO2is the major gas we want to talk about. By releasing CO2, we expect degassing to contribute to warmer climate on the long term and acidification of the oceans.
Hydrothermalism is a major process of chemical exchanges between the surface and the deep Earth: at high or low temperature, the mafic crust interacts with the ocean and large amounts of ions and metals are transfered from one reservoir to the other. The mafic crust is ultimately subducted and remixed within the convective mantle. At the bottom of the oceans, dissolution of mafic minerals at high temperature tends to buffer the composition of the ocean towards basic, releasing a lot of alkalinity. Precipitation of oceanic molecules into the crust tend to lower the carbonate content of the ocean and replenish the mantle in carbon and sodium.
Erosion of continental rocks is another process that transfer chemical elements from solid rocky reservoirs to the ocean. Alteration and erosion of rocks allows precipitation of carbonate and a transfer of alkalinity to the ocean. It is supposed to be today the most important source of alkalinity for the oceans. Also, most of the nutrients needed in the oceans come from the river fluxes like phosphorus. Such process could be compared to hydrothermalism in terms of effects, however the continental crust is felsic and the chemical supply from it should be different from that of the oceanic crust. And here, direct interaction with the atmosphere happens.
These processes are controlled by tectonics. The intensity of magmatism is controlled by
Xenon has radiogenic and fissiogenic isotopes. It is possible to measure the individual contribution of these nuclear reactions to the xenon budget as you see here for MORBs. 129Xe comes from the decay of 129I, whereas 136Xe comes from the fission of 244Pu and 238U. Each one of these decays has a specific half-life and the combination of the 3 gives an excellent view of the degassing processes that were dominating at the 10, 100 and 1Gy scales. the difference between the atmosphere and mantle rocks shows that the atmosphere has been extracted before the complete decay of 129I and 244Pu.
We can estimate the ratios between the radiogenic and fissiogenic xenon isotopes for a closed-system (a mantle that would have not degassed nor differentiated). That is what you see on a log scale (taking 136Xe coming from the fission of 238U as the reference). We can measure these ratios in present-day mantle rocks and we observe that 129Xe and 136Xe coming from 244Pu fission have been lost from the mantle more efficiently than 136Xe coming from 238U decay. Since 238U decay is long term, that means that degassing was much more efficient early in earh history. this has been already pointed out in several studies in the 80’s and more recently by yokochi and marty.
The goal of this work is to give a quantitative framework to interpret these data and first we have to know how Xe is degassed from the mantle throughout the history of the planet. Degassing happened in three main stages. The first is accretion: degassing occured through shock degassing and melting. The second, that can be more or less contemporary to the first is magma ocean formation and cooling: highly vigorous convection and gas release at the surface of the earth. when the magma ocean cooled sufficiently that solid-state convection could start, degassing occured through sub-surface magmatism. today degassing of the mantle is located in places where melts reach the surface like ridges, arcs and hotspots. In the first two stages, one can consider that degassing is complete and happens at a very short time scale. On the last stage, degassing of the mantle is partial since it occurs only in places where melts are generated, and at a slow rate.
To model degassing in the solid-state convection stage we can use a simple box model approach where the parent element (U, Pu or I) is simply decaying or extracted to crustal components, and where Xe is degassed. the flux of degassing D(t) is parameterized by the flux of mantle rock that melts, assuming that Xe is very incompatible so that all the Xenon in the zone that melts is transfered to the melt and hence to the atmosphere. the volume of mantle that melts depends on the surface area where melting occurs and the depth at which the solidus is reached. it is the intercept of the solidus curved given by petrology studies and the mantla adiabat that depends on the temperature of the mantle.
a probably more robust prediction of our model is the melting/processing history. in gray is the range of acceptable histories. to reproduce the values of the xenon ratios, the mantle has to be processed between 4 and 9 times. during the early cooling stage that corresponds more or less to the hadean, 75% of the processing occured. Hence most of the melting history of the Earth happened in the fisrt By of the planet. nothing is happening anymore in the present-day mantle.
Magmatic rates heat fluxes few times that of today…
our views of the early earth are always very imaginative since we still have a poor understanding of the physics to be used to describe the evolution in the first By of the planet. in this framework it is particularly diffiicult to assess the thermal state and evolution of the mantle during the early stages of our planet.
Indeed, the continental and oceanic crust in the Archean were significantly different than today. And these come from the radiogenic heating of the crust and mantle before 2.5Ga that was, at least, twice that of today.
Archean crust was at higher temperature, the rocks composing the crust become softer. As a consequence, they can flow upon the weight of a load or boundary stresses. Because a soft crust si not strong enough to support high elevations, modern-style mountains were forbidden during the Archean.
To test this assertion, we have built a 3D thin sheet model of the Archean lithosphere that we put in a sort of triaxial testing experiment. The column is subject to lateral stresses and we compute the thermal and mechanical response of the lithosphere. In these calculations, we vary the radiogenic heat content and we start from a steady-state geotherm. This assumption can be questionned in such ancient times.
Under these stresses, the lithosphere can thicken and laterally flow depending on its strength. When the Moho temperature exceeds 700°C, the deep crust is too soft and cannot support the mountain load. As a consequence, it flows under its own weight and the maximum plateau elevation that we obtain in our models is around 2000m. However, when the Moho temperature is closer to 500°C, the lithosphere is very strong and resists the gravitationnal force of the mountain building. As you see here, there is a smooth transition between Archean and Proterozoic style of mountain building. From crustal flow dominated to thickenning dominated. Therefore, the Archean hypsometry must have been much flater than today.
If we look at subsidece now, we have the same kind of regulation of the crustal thickness, that has been advocated by Bailey already. Here we show a 2D thermo-mechanical experiment in which we start from an isostatic equilibrium but what you see in black is an emplaced 6km thick flood basalt. The basalt is denser than the surrounding crust and the loading generates lower crustal flow for a hot crust. The time scale of this flow is short as you see here. To sum up: during the Archean the crust is much hotter and lower crustal flow is a dominating tectonic process that regulates crustal thickness very efficienty.
In the Archean the continents were also flooded. In a 1999 paper Nick Arndt showed that a lot of the Archean continental flood basalts were emplaced subaqueaously. Flatenning the continents allow easier flooding, but it is the thicker oceanic crust at that time that had the biggest contribution in flooding the continents.
Indeed, most of the studies tend to show that with a hotter mantle, the oceanic crust was much thicker, reaching probably 20km. Here you have observations of Proterozoic to Archean ophiolites supporting this idea, also developped in models of decompression melting. The crust is the light component in the oceanic lithosphere. Hence, thickenning the oceanic crust changes the isostatic balance. The topographic difference between continents and oceans was then smaller than today.
We computed such isostatic equilibrium, assuming hypsometric and bathymetric models based on our previous work and models of Labrosse and Jaupart (2007). Here is shown the distribution of altitude and sea-level with and Archean type setting.The mantle is 150°C hotter than today, the continents are smaller and flatter. In this calculation, the continents are under 500-800m of water and less than 5% of the surface of the Earth is emerged. Off course, the volume of the oceans may have changed with time and these calculations are indicative.
To have a feeling of what such changes of hypsometry and sea-level could affect the planet, I show you here such changes applied to the present day continental configuration. Most of the world would be under water exceprt the highest elevations generated in tectonic convergent settings.
