Seismologists have described the inner core as one of the last frontiers of global seismology (Tkalčić, 2015). The inner core is one of the most challenging regions of the Earth to image with seismic waves, making it a region of inherent uncertainty. The composition, structure, and temperature of the inner core affect the Earth’s heat budget, the formation of the magnetic field (geodynamo), and facilitate complex interactions between the mantle and core.
The center of the Earth has fascinated humanity for centuries, but the hypothesis of the inner core is a recent development. Edmond Halley was the first person to conceptually model the inner core; in 1686, he published a model where a solid sphere was separated from a spherical shell by a thick liquid layer. It was not until 1936 that Lehmann observed the Earth’s solid inner core from the arrival of P’ waves; the inner core was confirmed to be solid in 1971 (Dziewonski & Gilbert, 1971). We now know that Halley’s simplistic model of the Earth is accurate; the Earth is subdivided into four primary layers (crust, mantle, outer core, inner core), where the solid sphere (inner core) is separated by a liquid shell (outer core). The outer core is a liquid iron alloy with 5-10 % of light elements spanning 2900 km to 5100 km (4,400-6,000 K, 136-330 GPa). The inner core is a solid iron alloy with <5% of light elements spanning 5100-6400 km (5,000-7,000 K, 330-360 GPa) (Birch, 1952).
Recent seismological studies have provided constraints on the properties of the inner core. The most common technique is differences in P wave travel times; however, some studies have also employed free oscillations of normal modes and S wave travel times (PKJKP) to constrain the inner core.
Figure 1. (a) P-wave travel time curves within the inner core. (b) diagram detailing the ray paths through the Earth (figure from Deuss, 2014).
Differences in P wave travel times reveal that the inner core is: seismically anisotropic in the N-S direction, has lateral variation, and is composed of 2-3 layers.
Anisotropic in N-S direction
In 1983, Poupinet et al. analyzed PKIKP travel times for 400 different observations for five years. They observed a directional dependence; north-south stations were faster than east-west stations. This observation was later confirmed by Morelli et al. in 1986. They studied travel-time residuals of PKIKP between 170 and 180 degrees and proposed the inner core was anisotropic with cylindrical symmetry aligned with the axis of rotation; the N-S was observed to be 1% faster than E-W.
In 1986, Cormier and Choy analyzed differential travel times (PKP-Df, PKP-CD, and PKP-Cdiff) between 134-157 degrees, and they observed a 0.9% variation in the P velocity in the top 200 km. They inferred that the velocity variation resulted from structures at the base of the lower mantle or anisotropy in the inner core. In 1997, Tanaka and Hamaguchi confirmed their observation when they analyzed PKP(BC) -PKP(DF) in the east-west direction and observed heterogeneity in the top 100-500 km. They demonstrated that 43 degrees E-177 degrees E (quasi-eastern hemisphere) were regions of fastest P wave velocities, while 183 degrees W -43 degrees E (quasi-western hemisphere) were the slowest. Since then, the thickness of this heterogeneous layer has been debated, with 200-300 km representing a reasonable range (Iritani et al., 2014).
In 2002, Ishii and Dziewon´ski proposed the existence of an innermost inner core that is 300 km thick. Based on 325,000 travel time measurements from 1964-1994, they infer that the innermost inner core is transversely isotropic. This study was later repeated by Cao and Romanowicz (2007), who confirmed the existence of an innermost inner core but proposed the radius was 500 km thick.
In 1994, Kaneshim et al. analyzed differential travel times, PKPBC and PKPDF, and noticed a decrease in P wave speed with respect to PREM. They attributed this decrease to a small P wave gradient in the lower 300 km of the outer core and a large gradient at the top 300 km of the inner core (Kaneshima et al., 1994). Over time this measurement was refined (using PKIKP and PKiKP phases) to a steep gradient in the upper 80 km of the inner core and the lower 200 km of the outer core; this region is now referred to as the F region (Attanayake et al., 2014).
The combination of the F region with the two layers proposed by Ishii and Dziewonki makes up the three layers observed in the inner core. Wang and Song (2017) published a schematic shown in figure 2 highlighting our current understanding of the inner core. The first layer (F region) is about 80-100 km thick and is predicted to be isotropic. The second layer is the outer-inner core (OIC) that spans 100-700 km. This layer is fastest in the N-S direction, but the magnitude is heterogenous E-W. The third layer is the inner-inner core (IIC) which exhibits equatorial anisotropy, and the velocity is slowest at a 45-degree angle (Wang & Song, 2017).
Structure of the Inner core
Each of the three techniques above can constrain the inner core structure: normal modes and body waves were used to determine that the inner core is solid. Normal modes predict that the inner core is anisotropic and may possess dynamic topography that body wave observations support. PKJKP travel times predict the inner core is soft. While differences in P wave travel times predict that the inner core has lateral variation and is composed of 2-3 layers. Figure 2 from Wang and Song (2017) is a schematic of what we think the inner core looks like based on normal modes and body waves.
Figure 2. Model of a three-layer inner core. The first layer is inferred to be isotropic and shown in pink. The second layer, OIC (outer inner core), is anisotropic with lateral variation. The eastern region, in green, is where waves propagate faster relative to the western region, in yellow. The final layer, IIC (inner inner core), has the highest degree of anisotropy in the N-S vs. E-W direction. (figure from Wang and Song (2017)).
Understanding why the core looks this way is a key motivation for my research.
The cause of seismic anisotropy, lateral variations, and the formation of layers is heavily debated. The overall anisotropy is inferred to be the result of lattice preferred orientation of minerals. This lattice preferred orientation may arise from crystal flow, thermal convection, preferred growth orientation, or deformation from the magnetic field (Bergman, 1997; Buffett, 2009; Karato, 1999; Lincot et al., 2016; Romanowicz et al., 2016; Yoshida et al., 1996; Yukutake, 1998). The lateral variation is thought to arise due to deformation-induced texturing from Taylor column convection in the outer core (Frost et al., 2021). While the formation of the F region is thought to arise from a chemical exchange between the outer core and inner core (Cormier & Attanayake, 2013). The formation of the IIC and OIC may have formed due to a change in tectonic regime, the formation of a proto inner core, or a change in the geomagnetic field (Wang & Song, 2017).
I used experimental analogs to investigate how core materials deform under non-hydrostatic stress.
We are currently investigating Fe-Nitirides to improve the modeling of density and anisotropy of planetary cores and evaluate possible light element alloys as core components.
Anisotropy and Implications for the Earth
The inner core has implications for the Earth’s heat budget, the formation of the magnetic field (geodynamo), and complex interactions between the mantle and core. The geodynamo is powered by a thermal and chemical exchange between the outer and inner core. As the inner core grows, it releases heat (the latent heat of fusion) and separates a light element-rich phase; this, in addition to the temperature gradient in the outer core, powers the geodynamo and produces the magnetic field.
The formation of the inner core was paramount in forming the magnetic field, but its dynamic structure/topography is critical in understanding core-mantle interaction and the planet’s heat budget. Geodynamical studies have proposed that thermal heterogeneity in the mantle can be explained/explain the structure of the inner core (Aubert et al., 2008; Tkalčić, 2015). This is apparent in a geodynamical model by Aubert et al., 2008 (shown in figure 3); in the model, they consider the outer core flow to be indicated based on thermal anomalies in the mantle. The cold regions promote more heat flow from the outer core than the hot regions. This facilitates a strong heat flux on the inner core below cold regions forming faster crystal growth. The crystal growth speed impacts grain size, topography of the inner core, and localized anisotropy (Tkalčić, 2015). In Aubert et al., 2008’s model, the equatorial belt has faster growth than the remaining inner core. They infer this could form a distinct upper inner core topography over time (inner core grows at about 1 mm/yr).
Figure 3. Cartoon showing the predicted impact the mantle can have on inner core growth. The heat flux from the mantle to the inner is drawn as red and blue arrows. Cold regions of the mantle generate a large heat flux within the outer core; this heat flux promotes rapid crystallization and on the inner core. Warmer regions of the mantle result in a lower heat flux with the outer core resulting in slow crystal growth at the inner core (Figure copied from Taklcic 2015).