Magnetic Anomaly Calculator

A magnetic anomaly is what remains after the main field generated by the outer core is stripped from a measurement. The main field ranges from roughly 25,000 nT near the magnetic equator to about 65,000 nT near the poles, and the reference model subtracted from observations is normally the International Geomagnetic Reference Field. Crustal anomalies left behind are typically tens to a few thousand nanotesla, though banded iron formations and massive magnetite bodies can exceed 10,000 nT. Most crustal magnetization is induced, meaning the rock is magnetized by the present field in proportion to its susceptibility: sedimentary rocks sit around 0.0001 to 0.001 SI, basalt and gabbro around 0.01 to 0.1 SI, and magnetite-rich ore above 1 SI. Modeling a compact body as a uniformly magnetized sphere makes the anomaly amplitude scale with the cube of radius over depth, which is why depth dominates the response and why the half-width of the anomaly, about 0.77 times depth for a sphere, is a usable depth indicator. Several limitations follow from that geometry. Potential field inversion is inherently non-unique, so a small shallow body and a large deep one can produce identical surface anomalies, and depth estimates need independent constraint. The shape of an anomaly depends strongly on the inclination of the ambient field, so at mid-latitudes a simple body produces an asymmetric doublet with a negative lobe, which is why interpreters apply reduction to the pole before mapping. Remanent magnetization acquired when the rock formed can oppose the present field entirely; the Koenigsberger ratio comparing remanent to induced magnetization exceeds one in many volcanic rocks. Diurnal variation and magnetic storms must also be removed using a fixed base station.

Magnetic prospecting is older than most geophysical methods. Swedish miners were using dip needles to trace iron ore bodies by the seventeenth century, and in the 1870s Thalen and Tiberg built instruments in Sweden specifically for measuring field components over ore, turning the practice into a quantitative survey technique. The theoretical foundation had been laid by William Gilbert, whose De Magnete of 1600 argued that the Earth itself is a magnet, and by Carl Friedrich Gauss, who in 1838 applied spherical harmonic analysis to separate the field into internal and external sources, the same mathematical apparatus that reference field models still use. The airborne era began as a military project. During the Second World War the magnetic anomaly detector was developed to locate submerged submarines, with Victor Vacquier at Gulf Research contributing the fluxgate instrument around 1941. After 1945 the technology was redirected to geological mapping, making regional aeromagnetic coverage practical for the first time. Packard and Varian introduced the proton precession magnetometer in the early 1950s, and optically pumped alkali vapor instruments later pushed sensitivity to the 0.01 nT level. The most consequential application came in 1963, when Fred Vine and Drummond Matthews at Cambridge, and independently Lawrence Morley in Canada, interpreted the symmetric magnetic stripes flanking mid-ocean ridges as a record of geomagnetic reversals frozen into spreading seafloor. The International Association of Geomagnetism and Aeronomy issued the first International Geomagnetic Reference Field in 1968 to standardize the reference subtraction.