When Earth’s mantle melts, it doesn’t melt all at once. Instead, only a small fraction of the rock liquefies at any given time, typically between 0.1% and 20% depending on the setting. This process, called partial melting, is the origin of nearly all magma on Earth and ultimately builds both oceanic and continental crust. The specific way melting happens, what chemicals end up in the liquid, and what’s left behind in the solid rock all depend on where in the mantle the melting occurs and what triggers it.
Why the Mantle Doesn’t Melt Everywhere
The mantle is made of peridotite, a dense rock rich in iron, magnesium, and the mineral olivine. Despite temperatures reaching well over 1,000°C, most of the mantle stays solid because pressure increases with depth, and higher pressure raises the temperature needed for rock to melt. The boundary between solid and liquid for a given rock at a given pressure is called the solidus. For dry peridotite, the solidus starts around 1,120°C at Earth’s surface and climbs steeply with depth. The mantle only melts when something pushes conditions past that boundary: either the rock rises to lower pressures, something lowers the melting point, or extra heat arrives from below.
Three Ways the Mantle Melts
Decompression Melting
This is the most important melting mechanism on the planet. When mantle rock rises toward the surface, pressure drops, but the rock doesn’t lose heat fast enough to cool down. Its temperature stays nearly constant while the solidus temperature falls around it. Eventually the rock crosses the solidus and begins to melt. This happens most dramatically beneath mid-ocean ridges, where tectonic plates pull apart and mantle material wells up to fill the gap. The process has been described as arguably the most significant of all igneous processes, because it generates the magma that creates new ocean floor along roughly 65,000 kilometers of ridges worldwide.
Flux Melting
At subduction zones, where one tectonic plate dives beneath another, the mechanism is completely different. The descending slab carries water-rich minerals and carbonate sediments deep into the mantle. As the slab heats up, those minerals break down and release water and carbon dioxide into the overlying mantle rock. These volatiles dramatically lower the solidus temperature, allowing melting to begin at temperatures that would otherwise keep the rock solid. This is why volcanic arcs like the Andes and the Cascades sit above subduction zones: the mantle there isn’t hotter than usual, but the addition of water makes it melt anyway.
Hotspot Melting
Mantle plumes are columns of abnormally hot rock rising from deep in the mantle. They carry excess temperatures of 250°C or more above the surrounding mantle. That extra heat is enough to push peridotite past its solidus even in the middle of a tectonic plate, far from any ridge or subduction zone. Hawaii and Iceland are the classic examples. For plumes that also carry denser rock types like eclogite, even higher excess temperatures (around 550°C in the lower mantle) may be needed to stay buoyant enough to reach the surface. When a large plume does break through, it can generate enormous volumes of lava, forming features called large igneous provinces.
What the Melt Looks Like Chemically
The liquid that forms during partial melting is not a miniature copy of the original rock. Peridotite is ultramafic, meaning it’s very low in silica and very high in iron and magnesium. But when it begins to melt, the silica-rich components liquefy first because they have lower melting points. The result is a melt that’s significantly richer in silica than the source rock. At low degrees of melting beneath a mid-ocean ridge, this produces basaltic magma, which is mafic (moderate silica, still rich in iron and magnesium). If melting continues or the melt undergoes further processing in the crust, it can become intermediate or even silica-rich.
Trace elements behave in a particularly dramatic way during this process. Elements that don’t fit well into the crystal structure of mantle minerals, called incompatible elements, flood into the liquid phase even when only a tiny fraction of the rock has melted. Barium, lanthanum, and niobium, for instance, have partition coefficients below 0.01 for all major mantle minerals, meaning less than 1% of these elements stay behind in the solid. This concentrating effect is why volcanic rocks can be surprisingly enriched in rare elements compared to the mantle they came from.
What Stays Behind in the Solid
After melt has been extracted, the leftover solid is chemically and mineralogically different from what it was before. The original peridotite, known as lherzolite when it contains a full complement of minerals, progressively transforms into harzburgite as melting strips away its more fusible components. Harzburgite is depleted in silica, aluminum, calcium, and the trace elements that escaped into the melt. It’s essentially the skeletal remains of the original rock: tougher, more refractory, and harder to melt a second time.
The degree of depletion depends on how much melt was produced and how efficiently it drained away. In subduction settings where water flux is low, small amounts of melt can linger between mineral grains, causing local chemical redistribution without fully depleting the rock. Where water influx is higher, melting is more extensive and melt extraction more efficient, leaving behind a more thoroughly stripped residue. Studies of mantle rock samples brought to the surface as xenoliths (chunks carried up by volcanic eruptions) show that both lightly and heavily depleted rocks can coexist in the same region of mantle, reflecting this patchwork history.
How Much Melting Actually Occurs
The mantle rarely melts more than a small percentage at any one time. Beneath ocean basins, away from ridges and hotspots, the degree of partial melting in the low-velocity zone (a partially molten layer roughly 80 to 200 kilometers deep) is estimated at just 0.1% to 0.5%, based on how seismic waves slow down as they pass through. That’s enough to form thin films of melt coating the boundaries between mineral grains, but not enough to form large pools of magma.
Beneath mid-ocean ridges, melting is more substantial, typically reaching several percent as rock rises through a broad triangular melting zone. The roughly 7-kilometer-thick oceanic crust that forms at ridges worldwide is the product of this process. At hotspots with high excess temperatures, melting percentages can be higher still, producing thicker crust like Iceland’s or the massive basalt plateaus of flood basalt provinces. At subduction zones, the degree of melting varies depending on how much water is delivered by the sinking plate, but the water’s ability to lower the solidus means melting can begin at relatively modest mantle temperatures.
How Melt Escapes the Mantle
Once liquid forms between mineral grains, it needs to travel upward through tens of kilometers of solid rock to reach the surface. The driving force is buoyancy: basaltic melt is less dense than the surrounding peridotite, so it rises. At very low melt fractions, liquid forms an interconnected network along grain boundaries, allowing it to percolate slowly upward even without fractures. As melt accumulates, it can collect into larger channels and eventually into magma chambers in the crust.
The speed of this journey varies enormously. Percolation through grain boundaries is slow, on the order of centimeters to meters per year. But once magma enters fractures or conduits, it can accelerate dramatically. At volcanoes like Hekla in Iceland, magma appears to rise as a buoyancy-driven flow from deep storage zones to the surface, with the system continuously replenished from below so that the volume in the deep crust stays roughly constant even during eruptions. This plumbing system, stretching from mantle source to volcanic vent, is what connects the slow, diffuse process of mantle melting to the sudden, visible drama of an eruption.