The history of a landscape printed in a crystal...

What if we told you that we can read microscopic crystals to reveal the evolution of Earth’s landscapes through time? Although this may seem science fiction, this is possible thanks to modern advancements in low-temperature thermochronology! Tracking the evolution of modern landscapes in extraordinary detail via low-temperature thermocrhonometry is fundamental to quantify and discriminate between tectonic and climatic processes that drive erosion—a primary motivation of the COOLER project. Here, we provide an introductory overview to the science behind the COOLER project.

What is low-temperature thermochronology?
Thermochronologists investigate the thermal evolution of rocks within the crust or sediments within basins. The field of thermochronology, in part, revolves around the thermally activated diffusion of radiogenic particles (i.e., daughter products) produced during the radioactive decay of certain elements (parent nuclides) within a mineral’s crystal lattice. For instance, the radioactive decay of 238U (uranium-238) into 206Pb (lead-206) produces 8 atoms of radiogenic helium (4He), and this 4He will be retained within a given crystal only below a specific empirically derived temperature, known as the closure temperature. In apatite (a phosphate mineral), this closure temperature is about 75-85 °C; below this temperature, 4He particles are generally retained (we say that the system is closed), whereas at higher temperatures, 4He will have enough energy to escape the crystal lattice via diffusion (open system). If we know the production rate of such radiogenic daughter particles, we can measure the concentrations of these particles with that of their parent nuclide to calculate a thermochronological age—the date when this mineral cooled through its associated closure temperature.

Temperatures within the crust generally increase with depth, and geographically variable geothermal gradients define this relationship (such gradients are typically ~30°C/km). For thermochronological systems, this relates the closure temperature to a depth within the crust. For the example cited above, the depth of the closure temperature for helium on apatite ranges between about 2-3 km. With a thermochronological age, we are then able to trace back the location of minerals within the crust in the past, and infer rates of exhumation (i.e., their movement up towards the surface due to tectonics and/or surface erosion).

Thermochronologists have a suite of thermochronometers at their disposal. Each thermochronometer is defined by the isotopic system at play and the host mineral; the unique properties of the daughter element and the host mineral’s crystal lattice give each thermochronometer a specific temperature sensitivity. As such, numerous thermochronometers exist and can provide temporal constraints for a broad range of temperatures, from high temperatures (800-900 °C; e.g., zircon U-Pb) to low temperatures (30-70 °C; e.g., apatite (U-Th)/He) (Figure 1). This enables scientists to extract near complete thermal evolutions of Earth’s crust at different locations, and to in turn investigate rock exhumation pathways from deep within the crust and up to the surface.

Figure 1: List of common thermochronometers with their associated closure temperatures. Adapted and modified from Kronenberg et al. (2002).

What can low-temperature thermochronology be used for, and how does this relate to the COOLER project?
On Earth, the landscape is formed by the interplay between constructive forces that build relief, and destructive forces that reduce relief (Figure 2). Tectonics is the primary driver of relief; mountain ranges can be formed by contractional tectonic forces (e.g., the Himalaya), or even by extensional tectonic forces (e.g., the western U.S. Basin and Range province). These processes lead to surface uplift, which is counter-balanced by the main process that destroys rocks: erosion. Erosion modifies the surface of the earth and is primarily driven by the action of rivers, glaciers, gravity, and wind. These processes are modulated by climate through time, so that the shape of the landscape results from both tectonic and climate controls.

Tectonics and climate may interact through several mechanisms. One important interaction is based on chemical reactions: the weathering reactions that produce clays from silicate rocks consume CO2 from the atmosphere. Erosion also leads to the transport and storage of organic carbon (e.g. fallen trees, plants) into deep sedimentary basins. Both processes lead to cooling of the global climate on the timescale of millions of years. These mechanisms are commonly proposed as major drivers of long-term climate cooling since the start of the Cenozoic era (~55-50 Myr ago), broadly coinciding with the rise and onset of erosion within the largest modern mountain belt on Earth—the Himalayas. However, the consequence of Cenozoic cooling on global erosion rates, along with their ensuing dynamic feedbacks, remain poorly understood. A major challenge of the COOLER project is to quantitively assess the poorly understood physical nature between topographic relief and climate, and the impact this coupling has on erosion and weathering rates throughout the late Cenozoic, a period of dramatic shifts in Earth’s climate and a major reshaping of landscapes due to Plio-Pleistocene glacial-interglacial cycles. Conventional thermochronometric and modeling techniques lack the temporal and spatial resolutions necessary to adequately assess the relationships between topographic relief, climate, and erosion. The COOLER project aims to tackle these issues through (1) acquisition of high-resolution 4He/3He thermochronometric datasets from settings across the globe, and (2) next-generation thermo-kinematic/landscape evolution modeling.

Figure 2: Portion of the terrestrial crust illustrating the forces that shape the landscape and the concept of thermochronology. Here, contractional tectonic forces build mountain relief, which is also controlled by erosion modulated by climate. The interplay of these forces leads to rock uplift and exhumation. Here, we consider the helium thermochronological system in apatite (yellow mineral illustrated). The apatite cools through the PRZ (Partial Retention Zone) and closure temperature (Tc) and starts to record time. When the apatite reaches the surface, it can be sampled and analyzed to obtain its thermochronological age. Note that the resulting topography influences the temperature field in the crust where isotherms (lines of equal temperature) mimic the topographic relief.

Why are we interested in 4He and 3He in apatite?
As mentioned above, radiogenic helium (4He) in apatite is produced by the decay of 238U, 235U, and 232Th, with a minute contribution from the decay of 147Sm. Conventional apatite (U-Th)/He (AHe) methods generally involve single-step, high-temperature degassing and measurement of 4He from individual crystals, followed by parent nuclide concentration measurements (the U-Th-Sm component) to obtain a bulk mineral AHe date. Although conventional AHe dates provide invaluable insight into near-surface crustal cooling, single sample AHe dates are limited in the information they may provide. For example, Figure 3 reveals three contrasting thermal histories that will result in a single sample AHe age of ~1.7 Ma (from Shuster et al., 2005). If we want to address geologic problems that require resolving thermal histories at such high resolutions (e.g., ‘rapid’ vs ‘slow’ exhumation in Figure 3), we will need more information from this single sample. 4He/3He thermochronology provides an advantageous approach that allows measurement of the non-uniform spatial distribution of 4He within an individual crystal, which is highly dependent on the mineral’s thermal history for reasons explained below. 

Although thermochronologists often use the term “closure temperature” for simplicity, 4He actually begins to be partially retained within a broad specified temperature range referred to as the Partial Retention Zone (PRZ). The rate at which an apatite crystal cools through its PRZ will impact the retention, and thus distribution, of 4He within the crystal. For example, in a slow cooled setting, an apatite crystal will spend a relatively longer period of time within the 4He PRZ. This would allow more time for 4He near the grains edge to partially diffuse out, resulting in a 4He distribution that gradually increases towards the core (i.e., a rounded profile). In contrast, rapid cooling through the PRZ would result in more 4He along the grain edge and a less rounded concentration profile across the grain. Accordingly, the distribution of 4He within an apatite crystal can actually record the cooling history of its host rock—we just have to map it out.

Mapping out the spatial distribution of 4He within a crystal is not so simple, and requires a sophisticated analytical approach to achieve. Measuring the distribution of 4He requires an indirect measuring protocol that involves gradually heating a sample at increasing temperature steps, and measuring the extracted gas at each step. As mentioned above, natural 4He within U-bearing accessory minerals such as apatite is non-uniform. However, we cannot reveal the non-uniform distribution of 4He within a crystal unless we have a reference for uniformity—we must rely on tracer particles that are distributed uniformly within that same crystal. This is where the 3He comes into play. Naturally, minerals like apatite contain negligible concentrations of 3He, but we can actually synthesize 3He uniformly within a crystal at measurable concentrations via proton-irradiation. By exposing crystals to high-energy proton bombardment, spallation 3He can be uniformly induced within individual crystals following excitement of nearly all targeted nuclei within the crystal and surrounding medium.

Following sample irradiation, we are bestowed with individual crystals that now contain both a non-uniform radiogenic 4He distribution and a uniform synthetic 3He distribution. By measuring these gases simultaneously at multiple individual heating steps, we can map out the 4He/3He distributions within the crystal. Resulting 4He/3He profiles are reflective of the 4He distribution across the grain, and these ratio profiles are typically utilized to help constrain high-resolution thermal histories. This is perhaps best illustrated in Figure 3, which shows predicted 4He/3He profiles for three contrasting thermal histories that result in an identical AHe age. Simulated 4He/3He profiles may be compared with real measurements via inverse and/or forward modeling to determine thermal histories at that best satisfy observed data (e.g., the rapid exhumation scenario in Figure 3). 

Figure 3: Example from the Coast Mountains of British Columbia, Canada illustrating how single sample 4He/3He thermochronologic data can improve the resolution of thermal models derived from conventional apatite (U-TH)/He data. (a.) Three contrasting thermal histories that each result in a bulk apatite (U-Th)/He age of ~1.7 Ma. (b) Predicted and observed apatite 4He/3He ratio profile used to test three contrasting thermal histories. Measured data are most concordant with that expected for rapid exhumation (red lines). Adapted from Shuster et al. (2005).

How to measure the atoms of Helium?
Our research and projects are based on our capability to precisely and simultaneously measure low concentrations of both 4He and 3He within our samples. Their small atomic size (~10-10m) requires a specific technique called mass spectrometry. The base principle is to ionize our atoms of Helium with a high-energy electron beam, and then deflect the ions through a magnetic field based on their mass-to-charge ratios (m/z) (Figure 4). After precise deflection, the ions reach a detector where we can measure a signal proportional to the number of atoms hitting the detector.

The high-energy electron beam causes the He atoms to become electrically charged (ionized), typically by knocking off one electron out of the electron shield. These ions then pass-through a (variable) magnetic field that deflects each ion’s path to an extent that depends on both its mass and charge. This principle allows the selection of specific masses at the other side of the flight tube. For a specific magnetic field, we can thus focus He ions to precisely hit a metallic plate connected to a strong resistor (in our case 1012Ω). This metallic plate is called a Faraday cup (detector). As selected ions hit the Faraday cup, a current is discharged and its intensity can be measured (in Ampere). The value of the measured intensity is proportional to the number of ions composing the beam, and thus allows quantification of He concentrations in our samples.

In addition, for very small number of atoms (intensity greater than 10-16 A), we use another style of detector called SEM (Secondary Electron Multiplier). These detectors simply convert each ion that hits the metallic plate into several secondary electron emissions to virtually enhance the signal. Consequently, such amplification allows the detector to measure very small ion beams that would have been undetectable with the use of a Faraday cup.

Figure 4: Illustration of a typical mass spectrometer. On the left: The complete path of an atom of Helium travelling through the source, flight tube and magnetic field, and the collector. On the right: Illustration of the ionization and magnetic deflection. Once the He atom is injected into the machine and ionized by a stream of high-energy electrons, the resulting ions are accelerated (Vo) through parallel electric plates (Fe), and then deflected in a magnetic field (Fm) before they reach a detector. Image credit: left "Atomic Structure and Symbolism” by OpenStax, Rice University, Chemistry. Right: Zimmermann memoire 1999.

At the 4He/3He Thermochronology Laboratory at the University of Potsdam, we use a ThermoFisher Helix SFT mass spectrometer (Figure 5) using a Faraday cup for 4He signal and a SEM detector for 3He signal. The SFT is capable of multi-collection, which allows alignment of the 3He and 4He beams at the same time on both detectors. Simultaneous measurement significantly improves the precision on measurement for both elements.

Prior to inlet of He gas to the mass spectrometer, the aliquots are heated with a Diode Laser in an Ultra High Vacuum chamber to a set temperature. The gases from the sample are then purified with chemical traps called NP10 getters (to remove the actives species such as H2O or CO2). The He is then separated from other noble gases (Ne, Ar, Kr and Xe) with a cryogenic trap, where the He is trapped on a cooled charcoal at ~10K. This extraction and purification line (Figure 5) is necessary to ensure that only the He is analysed in our mass spectrometer.

Figure 5: On the left, the ThermoFisher Helix SFT with on foreground the Faraday cup and the SEM detector, the magnet, and the ionisation chamber at the back. On the right, the prep line with the getters to purify the gases.