Earth’s mantle is a 2,900-kilometer-thick layer of hot, dense rock sitting between the crust and the outer core. It accounts for about 67% of Earth’s total mass and roughly 84% of the planet’s volume, making it by far the largest layer of our planet. Though we often picture it as a uniform shell, the mantle is actually a series of distinct zones, each with different minerals, densities, and physical behaviors shaped by the extreme pressures and temperatures at depth.
What the Upper Mantle Is Made Of
The upper mantle extends from the base of the crust down to about 410 kilometers. Its dominant rock type is peridotite, a dense, coarse-grained rock that is rich in magnesium and iron. More than 50% of upper mantle peridotite consists of the mineral olivine, the greenish crystal you might recognize in its gem form as peridot. The rest is a mix of two types of pyroxene (orthopyroxene and clinopyroxene) plus an aluminum-bearing mineral that changes identity depending on depth.
At shallow depths, that aluminum-bearing mineral is plagioclase feldspar. Go deeper, and increasing pressure transforms it into spinel. Deeper still, it becomes garnet. These shifts aren’t just academic curiosities. They create subtle changes in rock density and seismic wave speed that scientists can detect from the surface, giving us a way to map the mantle’s internal structure without ever drilling into it.
The uppermost portion of the mantle, together with the crust, forms the rigid lithosphere, the tectonic plates that drift across Earth’s surface. Beneath that sits the asthenosphere, where temperatures are high enough that the rock behaves almost like a very slow-moving fluid over geological time. The minerals are the same, but the rock is softer and more pliable, which is what allows plate tectonics to work.
The Transition Zone: 410 to 660 Kilometers
Between 410 and 660 kilometers down, the mantle enters a region called the transition zone. The chemistry here doesn’t change dramatically, but the crystal structures of the minerals do. At the 410-kilometer boundary, the pressure is so intense that olivine can no longer hold its normal atomic arrangement. It collapses into a denser crystal structure called wadsleyite.
Around 520 kilometers, wadsleyite transforms again into an even more tightly packed mineral called ringwoodite. Both wadsleyite and ringwoodite have the same chemical ingredients as olivine (magnesium, iron, silicon, and oxygen), but their atoms are squeezed into progressively more compact configurations. These phase changes are what create the sharp seismic boundaries that geophysicists detect at 410 and 660 kilometers. The boundary at 520 kilometers is less consistent and appears to be sporadic rather than a clean global feature, which tells scientists something about how temperature and composition vary from place to place in the deep Earth.
One fascinating detail about the transition zone: ringwoodite can trap small amounts of water within its crystal structure. Laboratory experiments and a naturally occurring ringwoodite sample found inside a diamond suggest the transition zone may hold as much water as all of Earth’s surface oceans combined, locked away in mineral form rather than existing as liquid.
Lower Mantle Minerals
At the 660-kilometer discontinuity, ringwoodite breaks down under even greater pressure into two new minerals. The dominant one is bridgmanite, a magnesium-iron silicate that is likely the single most abundant mineral in the entire planet, making up a large fraction of everything between 660 kilometers and the core-mantle boundary at about 2,900 kilometers. The second mineral is ferropericlase, a simpler magnesium-iron oxide.
Together, bridgmanite and ferropericlase account for the bulk of the lower mantle. The density of mantle rock increases steadily with depth, ranging from about 3.3 grams per cubic centimeter near the top to around 5.4 grams per cubic centimeter near the base. That increase comes partly from these denser mineral phases and partly from the sheer weight of the rock above compressing everything below it. Temperatures in the lower mantle range from roughly 1,900°C near the top to over 3,500°C near the core boundary.
The D” Layer at the Mantle’s Base
The bottom 200 to 300 kilometers of the mantle, just above the molten iron outer core, is a strange and distinct region known as D” (pronounced “D double-prime”). Here, bridgmanite undergoes one final transformation into a mineral called post-perovskite. Research from Stanford University’s Extreme Environments Laboratory has shown that post-perovskite can absorb large amounts of iron, which dramatically changes its physical properties.
This iron enrichment helps explain some puzzling seismic observations in the D” layer, including zones where seismic waves slow down unexpectedly (called ultra-low velocity zones) and regions where waves traveling in different directions move at noticeably different speeds. The D” layer is also where mantle rock comes into direct contact with the liquid iron core, creating a thermal and chemical boundary that influences everything from how heat escapes the core to where mantle plumes originate.
Chemical Composition of the Whole Mantle
While the minerals change with depth, the overall chemistry of the mantle is surprisingly consistent. It is dominated by just four elements: oxygen, silicon, magnesium, and iron. These four make up the vast majority of every mantle mineral from olivine at the top to post-perovskite at the bottom. Smaller but important amounts of aluminum, calcium, and sodium round out the picture, and trace elements like chromium and nickel are scattered throughout.
The mantle’s composition is often described as “pyrolitic,” a model based on the idea that the mantle is roughly what you’d get if you mixed the basaltic rock that erupts at mid-ocean ridges back together with the depleted residue left behind after that melt was extracted. This model, while simplified, has held up remarkably well as a description of average mantle chemistry.
How Scientists Know What’s Down There
Nobody has drilled deeper than about 12 kilometers into the Earth, so direct sampling of the mantle is extremely limited. Instead, scientists piece together the picture from several lines of evidence.
The most important tool is seismology. Earthquakes send waves through the entire planet, and those waves speed up, slow down, or bend at boundaries where rock properties change. By analyzing thousands of earthquake recordings from stations around the world, geophysicists have mapped the mantle’s internal layering in detail. Every major boundary described above (410, 520, 660 kilometers, and the D” layer) was first identified through seismic wave behavior.
Physical samples do exist, though. Mantle xenoliths are chunks of mantle rock carried to the surface inside volcanic eruptions. These fragments, often just a few centimeters across, give scientists actual pieces of the upper mantle to analyze in the lab. Ophiolites are another source: these are slabs of oceanic crust and upper mantle that have been thrust onto continents through tectonic collisions. The Trinity ophiolite in northern California, for example, contains mantle peridotite samples including harzburgite and lherzolite that have been studied for their mineral chemistry and trace element patterns. However, researchers have found that xenoliths can be chemically altered by the very magmas that carry them upward, so interpreting their composition requires careful work to separate original mantle signatures from contamination.
High-pressure laboratory experiments fill in the rest. Scientists use diamond anvil cells and multi-anvil presses to squeeze tiny mineral samples to the pressures found hundreds of kilometers underground, recreating the conditions that produce wadsleyite, ringwoodite, bridgmanite, and post-perovskite. These experiments confirm what seismology predicts and reveal the precise pressures and temperatures at which each mineral transformation occurs.